the natural environment and the biogeochemical cycles

The Handbook of Environm.ental Chem.istry

Volume 1 Part B

Edited by O. Hutzinger

The Natural Environment and the Biogeochemical Cycles

With Contributions by H.-l Bolle, R. Fukai, l W. de Leeuw, S. W.F. van der Ploeg, T. Rosswall, P.A. Schenck, R. Söderlund, Y. Yokoyama, A.lB. Zehnder

With 84 Figures

Springer-Verlag Berlin Heidelberg GmbH 1982

Professor Dr. Otto Hutzinger

Laboratory of Environmental and Toxicological Chemistry University of Amsterdam, Nieuwe Achtergracht 166 Amsterdam, The Netherlands

ISBN 978-3-662-15324-6 ISBN 978-3-540-38597-4 (eBook) DOI 10.1007/978-3-540-38597-4

Library of Congress Cataloging in Publication Data Main entry under title: The Natural environment and the biogeochemical cycles. (The Handbook of environmental chemistry; v. I, pt. A-B). Includes bibliographies and index. l. Biogeochemical cycles. 2. Environmental chemistry. I. Craig, Peter, 1944-.11. Bolle, H.J. (Hans-JÜrgen). III. Series: Handbook ofenvironmental chemistry; v.l, pt. A-B. QD3l.H335 vol. I. pt. A, etc. [QH344] 80-16608 [574.5'222] AACR2

This work is subject to copyright. All rights are reserved, whether the whole or part of the material is concemed, specifically those of translation, reprinting, re-use of illustrations, broadcasting, reproduction by photocopying machine or similar means, and storage in data banks. Under § 54 ofthe German Copyright Law where copies are made for other than private use, a fee is payable to Verwertungs gesellschaft Wort, M unich.

© by Springer-Verlag Berlin Heidelberg 1982

Originally published Springer-Verlag Berlin Heidelberg New York in 1982.

Softcover reprint of the hardcover 1 st edition 1982 The use of registered names, trademarks, etc. in this publication does not imply, even in the absence of a specific statement, that such names are exempt from the relevant protective laws and regulations and therefore free for general use.

2152/3140-543210

Preface

Environmental Chemistry is a relatively young science. loteTest in this subject, however, is growing very rapidly and, although no agreement has been reached as yet about the exact content and limits of this interdisciplinary discipline, there appears to be increasing interest in seeing environmental topies which are based on chemistry embodied in this subject. One of the first objectives of Environmental Chemistry must be the study of the environment and of natural chemieal processes which occur in the environment. A major purpose of this series on Environmental Chemistry, therefore, is to present a reasonably uniform view of various aspects of the chemistry of the environment and chemical reactions occurring in the environment.

The industrial activities of man have given a new dimension to Environmental Chemistry. We have now synthesized and described over five million chemical compounds and chemical industry produces about hundred and fifty million tons of synthetic chemicals annually. We ship billions of tons of oil per year and through mining operations and other geophysieal modifications, large quantities of inorganic and organic materials are released from their natural deposits. Cities and metropolitan areas of up to 15 million inhabitants produce targe quantities of waste in relatively small and confined areas. Much of the chemical products and waste products of modern society are released into the environment either during production, storage, transport, use or ultimate disposal. These released materials participate in natural cycles and reactions and frequently lead to interference and disturbance of natural systems.

Environmental Chemistry is concerned with reactions in the environment. It is about distribution and equilibria between environmental compartments. It is about reactions, pathways, thermodynamics and kinetics. An important purpose of this Handbook is to aid understanding of the basic distribution and chemical reaction processes which occur in the environment.

Laws regulating toxie substances in various countries are designed to assess and control risk of chemicals to man and his environment. Science can contribute in two areas to this assessment; firstly in the area of toxicology and secondly in the area of chemical exposure. The available concentration ("environmental exposure concentration") depends on the fate of chemical compounds in the environment and thus their distribution and reaction behaviour in the environment. One very important contribution of Environmental Chemistry to the above mentioned toxic substances laws is to develop laboratory test

VI Preface

methods, or mathematical correlations and models that predict the environmental fate of new chemical compounds. The third purpose of this Handbook is to help in the basic understanding and development of such test methods and models.

The last explicit purpose of the Handbook is to present, in concise form, the most important properties relating to environmental chemistry and hazard assessment for the most important series of chemical compounds.

At the moment three volumes of the Handbook are planned. Volume 1 deals with the natural environment and the biogeochemical cycles therein, including some background information such as energetics and ecology. Volume 2 is concerned with reactions and processes in the environment and deals with physical factors such as transport and adsorption, and chemieal, photochemical and biochemical reactions in the environment, as weIl as some aspects of pharmacokinetics and metabolism within organisms. Volume 3 deals with anthropogenie compounds, their chemical backgrounds, production methods and information about their use, their environmental behaviour, analytical methodology and some important aspects of their toxic effects. The material for volume 1, 2 and 3 was each more than could easily be fitted into a single volume, and for this reason, as weIl as for the purpose of rapid publication of available manuscripts, all three volumes were divided in the parts A and B. Publisher and editor hope to keep materials of the volumes one to three up to date and to extend coverage in the subject areas by publishing further parts in the future. Readers are encouraged to offer suggestions and advice as to future editions of "The Handbook of Environmental Chemistry" .

Most chapters in the Handbook are written to a fairly advanced level and should be of interest to the graduate student and practising scientist. I also hope that the subject matter treated will be of interest to people outside chemistry and to scientists in industry as weIl as government and regulatory bodies. It would be very satisfying for me to see the books used as a basis for developing graduate courses on Environmental Chemistry.

Due to the breadth of the subject matter, it was not easy to edit this Handbook. Specialists had to be found in quite different areas of science who were willing to contribute a chapter within the prescribed schedule. It is with great satisfaction that I thank all 52 authors from 8 countries for their understanding and for devoting their time to this effort. Special thanks are due to Dr. F. Boschke of Springer for his advice and discussions throughout all stages of preparation of the Handbook. Mrs. A. Heinrich of Springer has significantly contributed to the technical development of the book through her conscientious and efficient work. Finally I like to thank my family, students and colleagues for being so patient with me during several critical phases of preparation for the Handbook, and to some colleagues and the secretaries for technical help.

I consider it a privilege to see my chosen subject grow. My interest in Environmental Chemistry dates back to my early college days in Vienna. I received significant impulses during my postdoctoral period at the University of California and my interest slowly developed during my time with the

Preface VII

National Research Council of Canada, before I could devote my full time to Environmental Chemistry, here in Amsterdam. I hope this Handbook may help deepen the interest of other scientists in this subjecL

o. Hutzinger

Contents

Basic Concepts of Ecology

S. W. F. van der Ploeg

Ecology: Some Definitions . . . . . . . . . The Science of Ecology . . . . . . . . . Organization Levels of Ecological Systems .

Species and Individuals. . . The Abiotic Environment The Biotic Environment . Limiting Factors . Adaptation . . . Habitat and Niche

Populations .. . . Introduction. . . Natality, Mortality and Dispersal. Dispersion. . . Limiting Factors Competition .

Communities . . . Introduction. . The Structure of Communities Species Diversity and Dominance. Communities Along Environmental Gradients

Ecosystems. . . . . Introduction. . . Trophic Structure. Production. . . . Biogeochemical Cycles.

Succession and Steady State. Succession. . . . . . . The Climax Concept . . Steady State and Stability in Ecological Systems

Major Ecosystems of the W orId Introduction. . . . . Terrestrial Ecosystems. . .

1 1 2 4 4 7 8

10 11 12 12 13 16 16 17 19 19 19 21 22 23 23 24 25 26 28 28 28 30 32 32 32

x

Freshwater Ecosystems . . . . . Marine Ecosystems. . . . . . .

Ecology and Environmental Problems Introduction . Pollution . . . . . . . . Exploitation. . . . . . . Environmental Disruption . Human Population Growth

References . . . . . . . . .

Natural Radionuclides in the Environment R. Fukai, Y. Y okoyama

Introduction . . . . Characteristics of Natural Radionuclides Classification of Natural Radionuclides .

Terrigenous Radionuclides . Cosmogenic Radionuclides .

Abundance in the Environment Radiation Effects . . . . . . Application of Geochemical Tracers

Transport Processes . . . Mixing Processes . . . . Sedimentation Processes . Exchange Processes . Pathway Indicators

References . . . . . .

The Nitrogen Cycles R. Söderlund, T. Rosswall

Introduction . . . . . Basic Chemical Considerations . . . . . . . . . . . . . . . . . . Chemical Transformations of Nitrogen Compounds in the Environment.

Introduction . . . . . . . . . . Nitrogen Fixation. . . . . . . . Mineralization and Immobilization Nitrification . . . . . . . . . . Denitrification and Nitrate Assimilation Abiotic Nitrogen Transformation .

Global Inventories of Nitrogen Introduction . . . . . . Atmospheric Inventories . The Aquatic System. . The Terrestrial System. .

Contents

36 37 40 40 40 41 42 42 43

47 47 50 50 52 53 56 57 57 58 58 59 59 59

61 62 62 62 62 65 65 66 67 68 68 69 70 71

Contents

Global Fluxes. . . . . The Ammonia Cycle The NOx Cycle ... The N 2/N 20 Cycle . Organic Nitrogen Transfers The Global Nitrogen Cycle.

References . . . . . . . . .

The Carbon Cyde A. J. B. Zehnder

XI

73 73 74 75 77 79

80

Introduction . . . . . . . . . . . . . . . 83 The Global Carbon Cycle. . . . . . . . . . 84

Carbon Balance in a Terrestrial Ecosystem . 86 Carbon Balance in the Ocean. . . 89

Photosynthesis . . . . . . . . . . . . . . 92 Photosynthetic Energy Conversion . . . . 93 Calcite Precipitation as a Result of Photosynthesis 97

The Carbon Dioxide Problem. 98 Sources of Carbon Dioxide. 99 Sinks of Carbon Dioxide . 99 Global Warming . . . . . 100 Environmental Responses to a Variation in Atmospheric Carbon Dioxide Content. . . . . . . . . . . 102

Biological Cycle of Carbon Dioxide 103

References . . . . . . . . . . . 106

Molecular Organic Geochemistry

P. A. Schenck, J. W. de Leeuw

Introduction . . . . . . . Normal Alkanes. . . . . . Acyclic Isoprenoid Hydrocarbons Steroids . . . . . . . . . . .

Occurrence and Diagenesis . . Steroids as Biological Markers

Triterpenoids . . . . . . . . . Occurrence and Diagenesis . . Triterpenoids as Biological and Maturation Markers.

Diterpenoids . . . . . . . . . . Polycyclic Aromatic Hydrocarbons. Epilogue .

References . . . . . . . . . . .

111 112 113 115 115 117 118 118 123 124 125 126

127

XII

Radiation and Energy Transport in the Earth Atmosphere System

H-J. Bolle

Introduction Nomenclature, Symbols and Units. . . . . . .

Structure of the Atmosphere and the Oceans . Radiation Terminology . . . . . . . . . .

Elementary Radiation Processes . . . . . . . . Relations Between Electromagnetic and Optical Properties of Matter Molecular Scattering . . . . . . . . . . . . . . . . . . . . . Deduction of the Rayleigh Scattering Coefficient and Phase Function Aerosol Scattering . . . . . . . . . . . . Representation of Aerosol Size Distributions. . . . . . Absorption . . . . . . . . . . . . . . . . . . . . Emission Under Thermodynamic Equilibrium Conditions Non-Thermal Emissions in the Upper Atmosphere

Atmospheric Radiation Field . . . . Solar Radiation . . . . . . . . Atmospheric Longwave Radiation Radiation Properties of Clouds. . Radiative Properties of Earth Surfaces Basis for the Theoretical Treatment of Radiative Transfer

General Energy Budget Equations for an Earth-Atmosphere System Energy Fluxes at the Top of the Atmosphere

Solar Irradiance . . . . . . . Planetary Albedo. . . . . . . Terrestrial Longwave Radiation Equilibrium Condition. . . . .

Energy Fluxes at the Earth Surface. Radiation Budget. . . . . . . Partitioning of Radiant Energy . Reat Flux into the Ground . . Flux of Sensible Reat into the Atmosphere Flux of Latent Real. . . . . . . . . . . Energy Used for Photosynthesis . . . . . Summary on the Partitioning of Energy at the Surface .

Energy Fluxes in the Atmosphere . . . . . Reat and Mechanical Energy Fluxes . . . . . . . . Deposition of Energy in the Atmosphere. . . . . . .

Energy Transports and Exchanges in the Atmosphere-Ocean System General Remarks. . . . . . . . . . . . . . . . . . Circulation Pattern in the Atmosphere and in the Oceans Reat Transport by the Oceans . . . . . . . . . . . . Energy Budget of the Earth-Atmosphere System . . . .

Contents

131 134 134 137 144 144 145 147 150 152 153 166 169 169 169 180 183 190 200 207 212 212 214 215 218 220 220 222 223 225 228 229 230 231 231 236 240 240 241 246 247

Effectsof Changes in the Concentration of Atmospheric Constituents on Energy Fluxes and Surface Temperatures . . . . . . . . . . . . . . . 252

Contents

Climate Research Aspects . . . . . . . . . . . . . Climatic Impacts of Specific Atmospheric Constituents.

Monitoring of Climate Parameters . Monitoring Strategy. . . . . . Baseline Stations . . . . . . . Upper Atmosphere Monitoring

List of Symbols . . . . . . . . . Frequently Used Numerical Values ~fureoc~ . . . . . . . . . . .

XIII

252 254 279 279 282 285 286 290 292

Subject Index . . . . . . . . . . . . . . . . . . . . . . . . . . . 305

Volume 2, Part B: Reactions and Processes

Basie Principles of Environmental Photochemistry. A. A. M. Roof Experimental Approaches to Environmental Photochemistry. R. G. Zepp Aquatic Photochemistry. A. A. M. Roof Microbial Transformation Kinetics of Organic Compounds. D. F. Paris,

W. C. Steen and L. A. Burns Hydrophobie Interactions in the Aquatic Environment. W. A. Bruggeman Interactions of Humic Substances with Environmental Chemieals.

G. G. Choudhry Complexing Effects on Behavior of Some Metals. K. A. Daum and L. W. Newland The Disposition and Metabolism of Environmental Chemieals by Mammalia.

D. V. Parke Pharmacokinetic Models. R. H. Reitz and P. J. Gehring

Volume 3, Part B: Anthropogenie Compounds

Lead. L. W. Newland and K. A. Daum Arsenic, Beryllium, Selenium and Vanadium. L. W. Newland Cl and C2 Halocarbons. C. R. Pearson Halogenated Aromatics. C. R. Pearson Volatile Aromatics. E. Merian and M. Zander Surfactants. K. J. Bock and H. Stache

List of Contributors

Prof. Dr. H.-J. Bolle Institut für Meteorologie und Geophysik Universität Innsbruck Schöpfstraße 41 A-6020 Innsbruck, Austria

Dr. R. Fukai International Laboratory of Marine Radioactivity IAEA Principality of Monaco

Prof. S. W. F. van der Ploeg Milletstraat 12 IV Amsterdam, The Netherlands

Dr. T. Rosswall SCOPE/UNEP International Nitrogen Unit Royal Swedish Academy of Sciences P.O.B. 50005 S-10405 Stockholm, Sweden

Prof. P. A. Schenck Dr. J. W. de Leeuw Dept. of Chemistry and Chemical Engineering Organic Geochemistry U nit Delft University of Technology Delft, The Netherlands

Dr. R. Söderlund Arrhenius Laboratory Dept. of Meteorology University of Stockholm S-10691 Stockholm, Sweden

Dr. Y. Yokoyama Centre des Faibles Radioactivites CNRS-CEA Gif-sur-Yvette, France

Dr. A. J. B. Zehnder Federal Institute for Water Resources and Water Pollution Control (EAWAG) Überlandstraße 133 CH-8600 Dübendorf, Switzerland



Milletstraat 12 IV Amsterdam, The Netherlands

Ecology: Some Definitions

The Science of Ecology

The world in which we live consists of living organisms and non-living structures. Often, relationships between organisms or between organisms and non-living structures are clearly visible. The science of ecology in its pure form studies the relationships of organisms with their environment. "Organisms" means allliving entities; this definition excludes relationships between non-living entities as a possible object of study for ecology. The term "environment" is meant in the sense of "the surrounding world," i.e., all entities, living or not, which surround a living entity. Thus for a grazing rabbit the environment includes for example other rabbits, grass, soil and weather.

Ecology is a study of relationships. These can be very complex or hardly recognizable. Therefore often studies are done on the relationship of one kind or organism, a species, with its environment. This type of ecology is called autecology. Even then reality mostly appears to be incredibly complex, as parts of the organism (organs or even cells) react differently to environmental influences. Hence ecophysiology has gained more and more importance, particularly in the last few decades.

The study of the relationships between groups of organisms and between socalled "communities" and ~he non-living (abiotic) environment is called synecology. On this level of complexity autecological issues are often neglected as these would render any understanding at the community level almost impossible.

Another division within the science of ecology can be made by discerning structural and functional aspects. In studying the structural relationships (e.g. the occurrence of various plant and animal species in a particular non-living environment), description of pattern and process is prevalent. In studying the functional aspects (e.g. the flow of energy from the sun through plants, herbivores and carnivores), measuring of flows and input-output relations is relevant. Oftenstruc-

2 S. W. F. van der Ploeg

tural ecology is rather descriptive, while functional ecology tends to be experimental.

Man is an organism. As such, his relationships with the environment are objects of study to the science of ecology. However, because of the importance of cultural aspects in the existence of Homo sapiens, human ecology is often seen as aseparate discipline. This does not imply that some basic concepts (as dealt with in later Sections) would not be applicable to our species.

In human ecology, the division into biotic ecology and social ecology is often used. Biotic ecology studies the reaction of human beings on environmental influences, particularly toxic substances, noise, radiation etc. Social ecology is concerned with the pattern and process ofhuman communities in relation to their environments, e.g. the use of communication systems like roads or the residential circumstances in a town as a result of structural processes. Basically human ecology can be viewed as a kind of autecology, be it that cultural aspects playa relatively important role. As human ecology makes use of theories and concepts from the social sciences while ecology requires contributions from physics, chemistry and earth sciences, the basic concepts of ecology can be regarded as the link between the natural and the social sciences [38].

Organization Levels of Ecological Systems

A system can be defined as an assembly of objects displaying some form of regular interaction or interdependence. Systems approach basically is a way of thinking about reality in which a collection of objects (or aseries of events) in considered to be a single entity. In ecology the systems approach is applicable because organisms always interact with other organisms and with the non-living environment.

Ecological systems are always open, i.e. there is an exchange (or input-output relation) of energy and matter with neighboring systems. The delimitation of ecological systems is thus often arbitrary: one could easily speak of sub- or supersystems, depending on the degree of complexity and the number of entities included.

Collier et al. distinguished four levels of organization in ecological systems, depending on the degree of complexity [5].

1. The Level of the Organism. On this level, ecological studies focus mainly at the relationship of the individual with its environment in the morphological, physiological or behavioral sense.

2. The Level of Populations. A population is a group of organisms of one speeies, living within a certain area. Such groups show group characteristics (e.g. density, distribution, age structure, rates of natality and mortality) which cannot be explained at the organism level. Next, populations interact with other populations and with the abiotic environment.

3. The Level of ( Biotic) Communities. A community is the assembly of different populations within a certain area. This combination of populations can be unique in space and time (or in pattern and process). Interactions between populations are very important for the composition of the community.

Basic Concepts of Ecology 3

4. The Level oj Ecosystems. An ecosystem is the system in which communities (including various populations) interact with the abiotic environment. At this level systems approach is most frequently used.

For each level of organization of ecological systems environmental abiotic factors are crucial. To a large extent they determine the possibilities for any lifeform to survive. Many influences like solar radiation, wind and rainfall exert their influence over several ecosystems. This stresses the open character of these systems, which is also revealed by migration of organisms. Sometimes one abiotic factor is dominant (e.g. solar radiation in the arctic regions), but often a specific local combination of abiotic factors, particularly soil and nutrients, allows for the existence of particular organisms, populations or communities. Therefore knowledge of abiotic factors is extremely important far the understanding of structure and function of ecological systems.

The definition of an ecosystem allows for considering the earth one ecosystem. This would, however, lead to a purely theoretical approach of properties of such a system. Therefore it is convenient to limit the extent of ecosystems to more easily recognizable units like a forest, a lake, or an estuary. Assemblies of such ecosystems can be called ecologicaljormations or major ecosystems. An easy division is the following: 1. Terrestrial ecological formations or biomes. These formations are largely de

fined according to climatic conditions. Examples are deserts, grasslands and forests, all existing in various forms on the continents.

2. Oceans and seas, both aquatic saline environments in which physical factors dominate (waves, tides, currents, temperature etc.).

3. Estuaries and seashores which are not only transition zones between land and water but also combine nutrient and energy inputs from both sides, thus displaying a rich variety of abiotic and biotic factors.

4. Freshwater formations, formed by inland water bodies. Examples are streams, rivers and lakes, again showing resemblances and differences from continent to continent. These ecological formations do not represent a fifth level of organization. In

some cases physical conditions can be the same for a spatial assembly of ecosystems but their relationships mostly take a general input-output form. Many formations are only recognized because structure and function of the ecosystems enclosed is roughly comparable (e.g. boreal forests).

The following four Seetions are devoted to the description of properties of the organismal, the population, the community and the ecosystem level of organization. Section 6 deals with changes in ecological systems in time and space, while Sect. 7 is devoted to abrief description of the major ecosystems of the world. Finally, Sect. 8 deals with the influence of mankind on ecological systems.

This introduction into ecology only deals very condensedly with some basic concepts. It must be stressed that any reader interested in a particular subject should obtain more detailed information from one of the existing textbooks of ecology, e.g. E. P. Odum (1971) [37]; Krebs (1972) [22]; Boughey (1973) [2]; Colinvaux (1973) [4]; Collier et al. (1973) [5]; McNaughton and Wolf(1973) [29]; Ricklefs (1973) [45]; Poole (1974) [43]; Whittaker (1975) [58]; Odum (1976) [38]; Ehrlich et al. (1977) [11] and Ricklefs (1978) [46].

4

Species and Individuals

The Abiotic Environment

Energy

s. W. F. van der Ploeg

Energy is a main source oflife, together with nutrients. The most important energy source for life on earth is, of course, the sun, but other energy inputs are cosmic radiation, the mo on (tides) and forces from the earth itself such as gravity and heat. Secondary sources of energy which are available to ecosystems are currents, waves, streams and wind. Ecological systems tend to use high-grade energy and to dissipate low-grade energy (heat), thus keeping the entropy within the system low and also operating under the laws of thermodynamics.

Green plants are able to combine CO2 and H20 into carbohydrates by absorbing light in pigment cells (containing chlorophyll):

6C02+12H20 hl 2.!~J >C6H1206+602+6H20. c orop y etc.

(from air) (to air)

These carbohydrates, in one or another form, constitute the living tissue or biomass of plants. However, not all energy fixed this way is retained. Plants also need energy for maintenance activities. This energy consumption is called respiration and can be generally represented as folIows:

C H ° +6 ° metabolie 6 CO +6 H ° +energy. 6 12 6 2 enzymes 2 2

(from air) (to air)

Thus accumulation ofbiomass in green plants (or net primary production) = energy fixed in photosynthesis - energy lost by respiration.

In bacterial photosynthesis, oxygen is not released. The reductant may be an inorganic compound (like H2S) or an organic compound. This type of photosynthesis particularly occurs in conditions unfavorable for green plants like tidal mudflats and H2S-rich stagnant lakes. Most green plants have a CO2 fixation via a C3

pentose phosphate cycle. However, in some plants (notably grasses) a C4 dicarboxylic acid cycle operates. C3 plants have their optimum in photosynthesis at moderate temperatures and light intensities, while C4 plants are favored by high temperatures and light intensities. C4 plants also need a smaller amount of water to produce the same biomass. Accordingly they dominate communities (deserts, grasslands ) in the subtropics.

Carbon dioxide and oxygen also limit photosynthetic processes. CO2 occurs in low concentrations (about 0.03 vol.-%) in the atmosphere. Increase of CO2 concentration causes increase in photosynthesis. High O2 concentrations inhibit the fixation of CO2 because most plants continue respiration in the light. This does not hold for C4 plants.

Animals cannot fix solar energy into living tissue. They are depending on already existing biomass and are therefore called heterotrophs, contrary to green plants which are called autotrophs. The biomass ingested consists of other organisms or particulate organic matter; this is converted to available nutrients by enzymes. Animals too, lose energy by respiration for maintenance activities. Endo-


thermic animals (mammals and birds) also spend much energy in keeping body temperature at a constant level. Movements require energy from almost all kinds of animals.

Most microorganisms gather their energy in the same way as the heterotrophs do, namely by breaking down matter by enzymes. Microorganisms living on dead matter (mostly organic) are called saprotrophs.

Other forms of energy also play a role in the maintenance activities of organisms. Obvious examples are the heat received from solar radiation, used for keeping body temperature at a steady level, and tidal movements which assure a continuous flow of nutrients for sessile organisms. Energy exchange mayaiso take place by convection, evaporation or conduction. In all cases the temperature ofthe organism will change.

Climate

Temperature dynamics on earth vary according to the orientation towards the sun. Thus heating and cooling differ from one place to another, causing atmospheric movements which are the basis for the worldwide c1imatic pattern. Organisms, however, are confronted with local c1imatic conditions which are the result of global c1imate dynamies. These local conditions are important in two ways: the regional c1imate and the microc1imate.

a) Regional Climate. Seasonal temperature, humidity (rainfall) and wind speed form the gross boundary conditions for organisms to survive. For example, oak (Quercus robur) flourishes in old inland dune valleys. I t can survive, in an adapted growth form, in relatively young dunes near the sea but it cannot withstand conditions in the outer dunes. Another good example form the plant species which find their optimum in humid, warm c1imates; they are not found in drier temperate regions (e.g. Dipterocarpacaeid trees of the tropics). Thus generally the distribution of species is limited to certain regional c1imatic conditions.

b) Microclimate. At the organismallevel, even regional c1imate is not decisive. Organisms have to interact with temperature, light, wind speed and humidity fluctuations in situ.

Organisms also respond to the particular combination of microc1imatic conditions rather than to each separate factor. For example, in dry conditions evaporation will increase with wind, while in humid conditions wind speed counteracts evaporation.

Particularly the combination of temperature and humidity is crucial for many organisms. For example, the mosquito Aedes aegypti changes behavior with increasing temperature and humidity, as is shown in Fig. 1.

Microc1imate also varies according to geomorphological conditions. For example, South facing slopes ofmountains or sand dunes receive a right-angle solar radiation which results in dry and hot conditions and also in severe fluctuations of temperature. Species occurring in these situations require a large physiological resilience. This holds both for organisms on the surface and for organisms inhabiting the upper soillayer.

6 s. W. F. van der Ploeg

Humidity

I I +--wb----fa----+" -c-

I I

Temperature 14 oe 33 oe

Fig.1. Activity of Aedes aegypti at different humidity and temperature conditions (After Haufe [17]). wb wing beating,Ja flight activity, c crawling (no wing action)

Matter

Anorganic and organic matter constitute the "food" for organisms. Nutrients may become available by means of physical processes (erosion, volcanic eruption, wind, currents). Both lack and abundance of nutrients are boundary conditions for all life forms.

In terrestrial ecosytems the soil is usually the most important nutrient source, particularly for plants and microorganisms. Soil in rivers, lakes and seas are also important both as substrate for life and as a source of nutrients which can be made available by stream or upwelling forces.

The combination of matter with energy results in both maintenance activities ("respiration") and accretion ofbiomass for any organism. Material gathering and use, however, is different for different species.

C, 0, H, and N are important building material. The role of C, 0, and H has already been partly discussed in relation to photosynthesis. N is, for example, extremely important in the composition of DNA or RNA structures and proteins. Other elements and their components which are of major importance are K, Ca, Mg, S, and P. Such nutrients are commonly called macronutrients because they are needed in large quantities.

Many other elements like Fe, Mn, Cu, Zn, Na, Mb are only required in small quantities. They are therefore called micronutrients. It should be noted that these elements are also essential in the metabolism of organisms, be it that the combinations and amounts needed are different from species to species.

A third general category ofmaterials important to organisms consists ofthe socalled non-essential elements, like Hg, Pb or Cd and their compounds. These have no obvious biological function but will nevertheless be taken up by organisms. In


Table 1. Nutrient quantities (ppm of dry weight) in different tissues of sugar maple (Acer saccharum) in deciduous forest. (After Whittaker [58])

Stern (bark) Branches Leaves (summer)

N 5,500 3,700 22,000 P 300 700 1,800 Ca 14,100 4,300 6,000 Mn 940 580 1,740 Fe 55 24 120 Cu 6 4 9

7

higher concentrations, these substances may be toxic to the organism as indeed many organic compounds are.

A fourth category of substances is formed by the allelochemies which are produced by one organism (or species) and are influencing other organisms (or speeies). More details will be given in the Section on the biotic environment.

To give an impression about the nutrient quantities in plant tissues, Table I shows some nutrient contents (in ppm of dry weight) in sugar maple in mixed deeiduous forests.

Materials may function, as a nutrient for example. However, materials also constitute the permanent abiotic conditions for an organism to live in. Air (N, 0, H, and carbon compounds) is surrounding any terrestrial organism. Water is the "physical" environment for aquatic organisms. Saline water bodies accomodate marine and estuarine organisms. Often large parts of such an environment are not really used (e.g. the 80% nitrogen ofthe atmosphere does not function for human respiration itself), but considerable changes in the composition of these "wrappings" may cause extinction of species.

Organisms, in reverse, influence the concentrations ofthe materials oftheir environment. BIue-green algae (Cyanophyta), far example, are often flourishing under eutrophie conditions (particularly high N and P concentrations) and release various toxic substances to the surface water. In terrestrial environments, the soil displays many of such interactions. In subarctic evergreen forests, for instance, the decomposing mat of spruce needles continuously releases organic acids which weather the parent material in the soil.

Tbe Biotic Environment

For an organism, many other organisms in the same area are meaningful. Functional relationships between organisms can be placed into several broad categories:

a) Structural Relationships. Plants often have a shelter function for other organisms. In forests many plants occur which are favored by the relatively stable physical environment (shade, humidity). Many animals, particularly terrestrial insects and other invertebrates, use vegetation structure as their "home." For parasites living in the intestines of animals, the host is essential both for food and for shelter. In lichens, algae and fungi are so intimately associated that the combina-


tion is considered a single species (this form of association is called mutualism, the association being necessary for the survival of both organisms).

b) Feeding Relationships. Animals eat plants or other animals, thus displaying a predator-prey-relationship (in a broad sense). Often species depend on particular plant or animal species or groups to feed on (e.g. all kinds ofparasitic organisms). Food searching behavior may be a very important maintenance activity. Lack of available "prey" organisms (e.g. due to extinction by Man) may result in failure to survive.

e) Chemieal Relationships. Organisms often release chemical substances which are attractive or repellent to other organisms. Chaparral shrubs like Salvia leueophylla, for example, release volatile terpenes which inhibit growth of grasses in the surroundings [34]. Such effects are called alleloehemieal or allelopathie effects. Animals like the monarch butterfly (Danaus plexippus) are distasteful to predators; other species apparently mimick the bright colors of such species. The unpalatability results from high concentrations ofterpenoids, alkaloids and similar chemicals. Released substances may be used both by organisms of the same species and by other organisms. [ps eonfusus, a beetle, releases a substance which attracts other individuals of the species to a tree which is suitable for infestation. The same substance also attracts a predator beetle species [58].

d) Behavioral Relationships. In social animals often some form of hierarchy is evident. This type of organization is useful in food gathering, defense and also in genetic selection. Large grazers (e.g. in the savannahs of Africa) also tend to display relationships among different species, which is particularly important in defense from predators and in the finding of food and water resources.

Territorial behavior within the species also confronts the organisms with the (im)possibilities for occupation of a certain area. This type of behavior is not limited to the own species; agressive male birds may chase away a number of individuals of other species. Reproductive behavior leads to genetic selection within the species. Differences in reproductive behavior mayaIso effect the forming of new species (speciation). In solitary animals, searching behavior of males is one of the most important activities.

The combination of these relationships is determining the possibilities for an organism or species to occupy a given area. In the next Section, interactions between populations for food and space will be dealt with in more detail.

Limiting Factors

The foregoing Sections have stressed the importance of the combination of abiotic and biotic influences for the organism or for the species. However, the absence of essential materials creates the lower limit for existence possibilities. In 1840 Liebig [26] formulated what is commonly named the "law of the minimum," basically meaning that "an organism is no stronger than the weakest link in its chain of ecological requirements" [37]. In relatively stable environmental conditions this is very relevant; if, however, the environment is unstable, all kinds of "minima" will frequently occur, to which sequence organisms often have no response at all.


Another complicating factor is the possibility of substitution; organisms may use a comparable element (e.g. strontium instead of calcium) to meet their requirements.

A second important concept regarding limiting factors has been indicated by Shelford in 1911 [48]. Organisms mayaiso have an ecological maximum in the supply of environmental factors. Thus organisms are supposed to have a range of tolerance for each environmental factor, the ends ofwhich represent the limits ofviability. For each factor this range will be different for different species (or even organisms). Commonly used classifications of organisms are stenohaline - euryhaline, stenothermal - eurythermal (referring to salinity and temperature, respectively), indicating a wide (eury-) or a narrow (steno-) tolerance range. However, often it is not possible to determine whether an organism occurs ne ar its optimum for a certain factor, as the influence of another (dominating) environmental factor may force it to live under sub-optimal circumstances. Certain orchids, for example, do very weIl in full sunlight, if they are kept cool; in nature they are only found in shaded areas [55].

It is extremely important to notice that the combination of limiting factors establishes the limits to survival for organisms or species. Here again, the factors creating absolute upper and lower limits of tolerance are determinitative. Hence, in studying the performance of individuals or species first attention should be paid to this kind of "controlling factors". F or example, in pollution ecology it is not only relevant to study the effect ofmercury, lead or cadmium on organisms; to know the real impact it is necessary to know the relative contribution of such substances to the performance of the species or community.

All environmental factors described earlier in this Section are potentiallimiting factors for organisms, as they constitute the essential basis for life. Some examples are:

a) Temperature Conditions. Although different organisms can stand different ranges of temperature (life in the desert or in the artic), their tolerance range is often narrow. Also variation in temperature (e.g. the day-night rhythm) is ecologically important.

b) Light. Photosynthesis, as already indicated, is related to solar irradiance. Figure 2 shows some responses of different organisms to increasing light intensity.

c) Water. Rainfall distribution throughout the year is a limiting factor to the geographical distribution of species. In micro-environments the humidity of the air or soil, partly due to rainfall, sets another limitation. Animals like springtails or spiders often cannot live at low relative humidity conditions.

d) Air. Concentrations of the main atmospheric gases CO 2 and O2 are often limiting factors. Again, in photosynthetic processes this is very clear. In aquatic environments such concentrations are also critical.

e) Nutrients. Macronutrients may become limiting because of high concentrations (e.g. eutrophication by P). Micronutrients (or trace elements) often set the lower limits for organisms. At least ten elements are definitely essential to plants [13]. They can be divided into three groups:

10

100

0/0 Photosynthesis

50

s. W.F. van der Ploeg

Light intensity

Fig.2. The relation between light intensities and photosynthesis for different organisms (After E. P. Odum [38]). 1 forest tree leaves, 2 com, 3 beach diatoms, 4 marine phytoplankton

(1) Mn, Fe, Cl, Zn, and V, for photosynthesis; (2) Mo, B, Co, Fe, for nitrogen metabolism; (3) Mn, B, Co, Su, and Si, for other metabolic activities.

f) SoUs. Texture and porosity of soils are important as to the availability of nutrients to plants and organisms. Poor drainage of soils results in reduced decomposition of matter, i.e. reduced release of nutrients. Lack of oxygen and accumulation of toxic substances may become limiting factors. Soil compaction (e.g. by trampling) inbibits plant root penetration.

Adaptation

The variability in nature is related to differences in environment. Comparable environments often contain the same species or species with comparable characteristics.

The presence of an organism in a certain area indicates that it is, in some way, adapted to the local biotic and abiotic environmental factors. Adaptation is a "fact" in biology. However, it is difficult to make out whether an organism has been adapted to its environment or has found a suitable environment or has made the environment more suitable. Adaptation may be genotypic which means that there is a change in the genetic structure of an organism (often called natural selection) or phenotypic which means that there are non-hereditary changes. The combination of both types of adaptation is rather common. Evolution is often a long-term process; therefore genetic adaptation to a sudden change in environmental factors is not likely. If the change of the environmental factor still falls within the limits of tolerance of the organism, phenotypical adaptation may be possible.

A biogeographical expression ofthe adaptation concept is Bergmann's rule: the size of homoiothermic animals in a c10sely related evolutionary line increases along a gradient from warm to cold temperatures. This is probably demonstrated by deer, rabbit and fox species, and also by man. The explanation for tbis difference lies in surface to volume ratios. The Bergmann rule, however, is certainly not accepted fully by all ecologists.


Convergence, i.e. a similarity in morphological and physiological characteristics of organisms occurring in similar climates, can be demonstrated with the difference between C3 and C4 plant species. Another example is formed by the "life-form spectrum" of Raunkiaer [44], which includes the following major groups of plants:

- Phanerophytes: trees and shrubs; buds are hardly protected; common in the tro-pics

- Chamaephytes: buds are at ground level during "difficult" seasons; common further away from the tropics

'--- Hemicryptophytes: buds (in dormant form) in the upper soil; mainly in temperate regions

- Cryptophytes: buds deeply buried; mainly in regions with extreme climatic conditions.

Ecotypes are populations within a species that show differences which are related to differences in environment. Ecotypes are partly genetically different, often due to geographicalor ecological isolation. MeNeilly [30], working on copper tolerance of the grass Agrostis tenuis on and around copper mines, found that the change from tolerance to non-tolerance of the plants occurred within about one meter. Away from the mine also tolerant individuals were found. MeNeillyeoncluded that selection pressure on toxie soils is strong and is in favor of tolerant genotypes, while selection pressure on non-toxic soils is weaker but relatively favors non-tolerant genotypes. Another example is the three-spined stiekleback (Gasterosteus aculeatus) which eonsists of a marine "race" , in Europe oeeurring in seas from the aretie through Holland and Belgium, and a freshwater "race" extending from Belgium to the Mediterranean region. The two raees are effeetive1y species because they are eeologieally isolated by the faetors temperature and salinity.

Habitat and Niche

The habitat of an organism is the place where it lives. This place may be deseribed in geomorphological, climatologieal and (sometimes) vegetational terms. Habitat is, in fact, the "address" of an organism.

The niche whieh is oceupied by an organism eould be ealled the "profession" of that organism, referring to the aetivities of it. This "profession", or funetional role, is determined by habitat properties, by food re1ationships and by its tolerance towards environmental faetors. Actually, defining the niche of an organism would require a definition of all biotic and abiotie faetors an organism is dealing with. U sually three types of niehes are designated:

a) The spatial niche, formulated by Grinnell in 1917 [15]. This type ofniehe is also called the habitat niehe and refers to the spatial properties of the environment.

b) The trophic niche, formulated by Elton in 1927 [12]. This type of niehe refers to the functional status of the organism towards other biotie elements of the eommunity.

e) The multidimensional or hypervolume niche. This niehe type, suggested by Hutchinson in 1957 [19], refers to the different environmental factors deeisive for the existence of an organism. Each faetor adds a niehe dimension.


Hutchinson also formulated the fundamental niche in which competition with other organisms is excluded, and the realized niche which is occupied under biotic constraints.

The importance of the concept of the niche is that it splits up the environment into small structural/functional parts which are occupied by species (or by certain populations of species). This type of splitting up renders the possibility of analysing the components of a community as regards structure and function, without ignoring the fact that the community itself exists. Therefore the concept of the niche is dealt with in textbooks often in the community Chapters. In some cases, niches are discussed at the population level, mainly because interspecific and intraspecific competition is believed to result into occupation of niches.

It is clear that the concept of the niche is dynamic. Changes in biotic or abiotic conditions can be regarded changes in niche composition. The organism or species has to adapt itself to these changes, or to move or die. Adaptation itself mayaIso be regarded as change of niches.

Fortunately, the ecological time-scale rather counts in decades or centuries than in months or years. Therefore the niche can usually be regarded as a fair ecological description of the role of any species. Pollution and other environmental degradation, however, may cause such perturbations that the niche arrangement of a specific community may be changed completely in a short time. In such cases, only species with wide tolerance ranges will be able to occupy the newly formed niches.

Populations

Introduction

A population has been defined as a group of organisms of one species living within a certain area. A species thus consists of several populations; in one area populations of different species live together.

Population ecology is probably the most developed branch of ecology. This is mainly due to the fact that it is often difficult to study all populations of one species (because ofthe sometimes worldwide distribution) and to the problems in studying communities. However, even populations are not as easily recognizable as it seems: intraspecific variation may be partly or completely genetically bound, resulting in neighboring populations of one species with different genetical characteristics.

Organisms are born and die. For a population this means that there are rates of natality and mortality which are partly dependent on the genetic information ofthe species itself, partlyon external factors. Population dynamics, which also include the dispersal of organisms, are an important branch in population ecology.

Organisms of a species usually do not live scattered all over the earth. In order to survive as a species it is necessary for males and females to meet. Therefore many kinds of aggregations of individuals occur, ranging from several pairs of golden eagles in the whole of Scotland to thousands of puffins on a rocky shore. These density and distribution patterns are also an important object of study in population ecology.


Finally, populations show relationships with each other. Predation and parasitism are limiting factors for populations. Competition for resources within and between the species leads to exclusion of populations and species from an area.

In population ecology the number of models trying to describe population processes has drastically increased during the last few decades. In this Seetion, however, not much attention will be paid to this quantitative and modelling work as it is the specialist's realm and is not necessary for a general insight into the concepts of population ecology.

Natality, Mortality, and Dispersal

Existing populations have a density, i.e. the population size in relation to a unit of space. Usually density is expressed in numbers ofbiomass per unit area or volume. Density may range from several individuals per 100 square miles to thousands per millilitre. Densities are dynamic quantities; they vary with time because populations grow or decline.

Populations also show an age distribution, being the fraction of the population falling in successive age classes. This age distribution is important because of the relation with reproduction. The so-called age pyramids show this age distribution within a population; an example is given in Fig.3.

22

Ag e 20

(monthsl

1

5 10 15 20 25 % of population

Fig.3. Age pyramid of the vole Microtus agrestis. (After Leslie and Ranson [25])


The production of offspring in a population is called natality. The natality rate can be expressed as the number of offspring produced per unit population size per unit time. Another characterization is the age specijic birth rate, being the average number of female offspring produced per unit time by a female of a specific age class of the population. Natality rates greatly differ from species to species: fish may produce millions of eggs per year, elephants only produce one young in two years at a maximum. Natality is never maximal because environmental conditions act as limiting factors. Food resources are obviously potentially limiting. For example, the production of eggs by the spider Erigone arctica has been shown to decrease if the population size of the main food resource, a springtail (Hypogastrura viatica) , declines [59]. Another limitation is set by weather conditions: many birds produce less offspring in bad summers.

M ortality can be expressed as the number of individuals dying per unit of time. The age specijic death rate is the number of individuals of a certain age class dying in a (short) period divided by the number of the same age class that were alive at the start of that period. Organisms may die of "senescence", which can be called physiologicallongevity. However, it is more important to consider the ecological longevity, i.e. the (empirical) average longevity of the individuals of a population under specific environmental conditions. Usually the ecologicallongevity is much shorter, for instance because of predation or disease.

For specific age classes (also called cohorts) survivorship curves can be constructed, showing the diminishment of such a class through time. An example of such a curve is shown in Fig. 4. Next, life tables can be constructed which summarize population statistics. Table 2 gives the numbers corresponding to Fig.4.

Natality and mortality together constitute the growth rate of the population. Under stationary age distributions and environmental conditions the specijic growth rate (i.e. the population growth per individual) is constant and the population grows exponentially. For this constant usually the symbol "r" is used. The equation for growth rate is

1000 Numbers

800

600

400

200

dN =rN dt

15 25 35 45 55 Age in days

(N =numbers, t=time).

Table 2. Life table for adult honeybee workers. (After Sakagami and Fukuda [47])

Days

5 15 25 35 45 55

Numbers living

1,000 962 912 795 551 44

Numbers dying

38 50

127 244 507 44

Fig.4. Survivorship curve ror adult honeybee workers. (After Sakagami and Fukuda [47])


VI c: QJ

Cl

\

A

K- - - - - - - - -

\

B

Time

Fig.5. J-shaped A and S-shaped B growth fonns of populations. For explanation see text

15

Actually r is the difference between instantaneous birth and death rate. r is usually called the intrinsic rate o/natural increase. Within a species it may vary according to age distributions and specific reproduction capacities of age classes. Values of r have been calculated for many species. For instance, under optimallaboratory conditions the brown rat (Rattus norvegicus) has an r of 0.1 04 on a week base; the mean 1ength of a generation is 31.1 weeks, resulting in a doubling time for the population of 6.76 weeks [24].

Population increase often takes a very marked form. At least two basic patterns can be designated: the J-shaped growth form and the S-shaped (sigmoid) growth form. In Fig.5 both growth forms are shown.

In the J-shaped form the population increases exponentially and then is stopped abruptly by any factor becoming limiting. In the S-form the initial increase is slow (lag-phase), then rapid (log-phase) but slows down as environmenta1 factors become more and more limiting. In the latter case the following model is adequate:

dN _ N(K-N) dt -r K '

in which "K", the upper level asymptote, is called the carrying capacity. The life history (or 1ife cycle) of a species gives a detailed information about the

way life is being organized from birth to death, with emphasis on reproduction strategies. U nder pressure of the circumstances, the life cycle may change considerably, resulting in different reproduction strategies and, consequently, another rate of increase (r). Evolutionary selection favoring high r values is called "r selection" [28]; it is usually shown by so-called colonizing species which are associated with unstable or newly formed habitats. "K selection", on the contrary, favors optimal use of available resources and maximum prob ability of survival for the individual itself and its offspring. In this case r values are relatively low.

Population sizes often fluctuate with time. In the case of K selection species, densities vary around the carrying capacity level. In populations with a J-shaped growth type fluctuations may be violent. Such fluctuations may be inherent to the species itself or may be caused by changes in environmental conditions. Of course there is often also a seasonal variation, e.g. due to climatic influences.

Population densities also change because of dispersal, i.e. the immigration into and emigration from a population by individuals. Dispersal constantly occurs on a small scale because individuals enter or leave the population. However, mass emi-

16 s. W.F. van der Ploeg

gration or immigration may completely change the balance between the population and its environment. Particularly r selection species have a high dispersal power (calIed vagility), thus being able to emigrate from overcrowded areas. In this way dispersal supplements natality and mortality in the overall population density.

Dispersal may take place on an annual or a perennial base. Many spider species have been reported to show aeronautic behavior, at least in their immature stages, thus being able to cover long distances. This behavior occurs each year and has a clear density-reducing function. European starlings, on the other hand, after being introduced into the US in 1890 in New Y ork, had spread from there up to the West Coast by 1949 [20]. FinaIly, dispersal is also important because it effects gene exchange between populations.

Dispersion

The way members of a population are distributed through space is called the pattern of dispersion. Three general patterns are recognizable: random, uniform or clumped. Random distribution seldom occurs; only in homogeneous environments there may be no stimulus for individuals to aggregate. Uniform distribution may occur in situations where (sessile) animals or plants would outcompete each other if they were arranged in a clustered way. However, the large scale pattern of such a distribution is often of a clumped form.

Clustering is most usually found in populations. This clustering can again take the three forms which were described for dispersion. Of these forms, aggregated clustering is most common.

Dispersion can have different causes. In animals requiring a high relative humidity (e.g. springtails), clustering takes place on wet places if the environment is drying out. Low temperatures or high wind speeds urge all kinds of arthropods to stay preferably in the lee side of a grass tussock. Thus dispersion may have a nonpermanent character.

Another useful division of dispersion patterns was made by Pielou [41]. Two types of dispersion are characterized by the words intensity and grain. Intensity refers to the number of individuals occupying a certain part of the area, while grain refers to the size of that part of the area.

Aggregations of individuals may be beneficial in terms of survival. For instance, groups of fish were more tolerant towards poisonous substances than isolated individuals were, mainly because of secretions counteracting the poison [1]. Therefore undercrowding (as weIl as overcrowding) may be limiting.

Limiting Factors

Populations are being limited in growth and survival possibilities by a vast range of factors. The factor having the most important influence is often called the key factor. Factors may be dependent or independent of the size of the population. In the case of density-dependency the population is affected in its growth rate by its own density. This type of limitation has been observed in many species (e.g. the spruce budworm, see Fig.6). However, in many cases it is not clear which factor exactly causes this dependency.


.5

Survival rate

10

17

100 1,000 10,000 Numbers

Fig.6. Relationship between survival of smalliarvae and population growth of the spruce budworm. (After Morris [33])

Food relationships may cause density-dependent changes in the population. If food shortage occurs in the habitat of a population because of excessive population growth, population size will decline sooner or later.

Food relationships also include predator-prey relationships, or, in general, exploitation interactions. Herbivore animals may regulate plant species populations while carnivores regulate the herbivores. In the case of continuous generations simple mathematical models, as developed by Lotka [27] and Volterra [52] show a regular oscillation of predator and prey numbers. In laboratory experiments this regularity is difficult to test, as was already shown by Gause [14] in his famous study on the protozoans Paramecium caudatum and Didinium nasutum. In the field it is often not clear whether a predator is actually regulating the prey population. Predators may answer to an increase in prey density in two ways: (I) a numerical response which means an increase in the predator density, (2) a functional response in which the consumption pattern of predators changes.

An example of these responses is shown in Fig. 7. Another type of species interaetion possibly leading to limitation is parasitism in which one population affeets the other one, being dependent on that population.

Abiotic faetors mayaiso limit population size. However, this possibility is still an unsolved problem in population eeology. Most ecologists eonsider food shortage the most eritieal faetor in the regulation of population size. Nevertheless elimatie eonditions (temperature, moisture) and physieal or ehemieal factors un. doubtedly restriet the distribution of populations, as has already been indicated earlier. Once a population has become established, adverse conditions may temporarily reduce the density but usually the population can quickly recover.

Competition

Competition can be defined as the interaction between individual organisms which use a resource that is in short supply. Resources include food (or nutrients), space, shelter ete. As populations normally co-exist in space, competition may be considered one of the most important regulation meehanisms of populations.

18

QI '-u

~ ..'1 ro E E ro E

..... 0

VI '-

25

15 /

;"

I

x x ............. _._. (3) A

,,-./ x

/

QI ..CI E ::J

Z

o~/ _____ -cc>----<~<>-- (1)

5~ i-.;.~ ~ -. ---A - ~ - - - - - ( 2 )

~ >.400 e: ro QI "0 c- '-o QI 300 VI c-e: o o u o u

ro E E 200 ro

..... E o '-o ~ 100 z

/ I

I I

I

__ A ______ (2) .- A

/

A '"

"'lt A

I

~ ..,.~x I ~?-o

B

o (1 )

2 4 6 8 10 12

S.W.F. van der Ploeg

Numbers of cocoons/acre(x105)

Fig.7. Numerical A and functional B responses of a deer mouse (Peromyscus 1) and two shrews (Blarina 2 and Sorex 3) to different densities of the cocoon of a sawfly (Neodiprion). (After Holling [18])

Intraspecijic eompetition me ans the interaetion between individuals of the same population in using searee resourees. This ean take the form of a scramble in whieh all individuals try to get their pieee of the meal, or in the form of a contest in whieh behavioral interaetion leads to the use of the resouree by the "winner". Territorial behavior is thought to be a clear manifestation of eontest eompetition. If the density of a nesting bird population beeomes high, some individuals will fail to breed at all beeause of short supply of territories. Thus this eompetition limits population growth at least.

Interspecijic eompetition is defined as the interaetion between organisms striving for the same searee resouree. Darwin (1859) already suggested [7] that speeies similar in behavior showed the most severe struggle for existenee. In his laboratory study on two Paramecium speeies, Gause [14] found that under eo-existing eireumstanees one speeies was eliminated while it eould easily survive if not in eompetition. Hardin [16] formulated the competitive exclusion principle. Shortly said this principle states that perfect competitors cannot co-exist in the same time-spaee dimension. This implies that if two eeologieally similar populations live together in


one area, there must be some difference at least. The concept ofthe niche is strongly related to this principle. In field studies the competitive exc1usion principle has been rarely demonstrated. The studies of Lack on cormorants and shags [23] which appeared to feed and nest differently, and of Connell [6] on barnac1es which are differently zonated on a rocky shore, are two of few examples. Actually the principle is almost impossible to prove as it is difficult to understand all relationships of field populations with the surrounding environment. Pontin [42] suggested that colonies oftwo ant species (Lasius niger and L.flavus) may indefinitely co-exist because intraspecific competition may be far more important than interspecific interactions.

Another possibility is to regard the exc1usion principle as a cause of genetic change of populations. Such a change is often called ecological isolation which is thought to result, in due time, in the forming of new species with different requirements. Spatial isolation, the effect of which is called allopatry leads to similarity or convergence in habitat and resource use in different geographical regions. On the other hand, sympatry (two c10sely related species living in the same region) causes divergence of species characteristics. There are also temporal isolation mechanisms, e.g. differences in diurnal or seasonal activity, and morphological, physiological or behavioral differentiations of resource exploitation mechanisms.

Communities

Introduction

A biotic community is any assemblage of populations within a specific area. Size is not important; the extent may vary from a hollow tree to an ocean. In a community populations of species are related to each other in various ways and a community is therefore more than a simple addition ofpopulations. Next, communities differ in their dependency on neighboring communities. This is partly related to the spatial and temporal extent of the community.

At the community level of organization several issues are important. The structure of the community, consisting of growth forms, stratification and horizontal pattern is often very characteristic. The same holds for differentiation in time. The variety of species is another important feature of communities. Within this diversity differences in ecological dominance of species occur. Finally, communities vary along environmental gradients; in such cases the composition of the community is regarded in strict relation to major environmental factors.

The above characteristics of communities regard structural properties. Functional properties of communities will be discussed in the Section on Ecosystems in relation with flows of energy and matter.

The Structure of Communities

The morphology of communities can be called physiognomy. The physiognomy is not always c1early visible, e.g. in the case of microscopic plankton communities. On the otherhand, coral reefs and particularly terrestrial communities display very characteristic physiognomies.


Growth forms are important for the structure of communities. For plants such forms are constituted by height, stern and leaf form, deciduousness or evergreenness etc.

A general classification is the following: - Trees (larger woody plants) - Lianas (woody climbers) - Shrubs (smaller woody plants) - Epiphytes (plants growing on other plants, above ground) - Herbs (plants without woody parts above ground) - Thallophytes (lichens, mosses). Another classification is according to life forms which has already been given earlier.

The growth forms also result in a more general vertical structure of terrestrial communities, called stratification. Actually plants within a community are grouped along a vertical gradient of light intensity. Complex forests consist of: - a canopy, formed by the tallest trees - a lower tree layer, of smaller or younger trees - a shrub layer, receiving only 5-10% of the sunlight - a herb layer, possibly accompanied by a moss layer, receiving 1-5% ofthe sun-

light. Animals are also stratified vertically, regardless whether they are very mobile

or not. Birds nest in alllayers, arthropod species differ from the canopy to the soil surface and in the litter layer of the top soil complete communities of cryptozoic animals (invertebrates) exist.

In aquatic communities there is also a clear stratification. A lake can be divided into three zones (see Fig.8):

Land Lake

LlT: LlM

'-- --- --- ----LB

PRO

Fig.8. Stratification of a lake. LIT=Littoral zone, LIM=Limnetic zone, PRO=Profundal zone, LB= Light boundary. For explanation see text

- the profundal zone (bottom and deep water, almost no light penetration) with mostly organisms attached or resting on the bottom (benthos);

- the limnetic zone (open water with light penetration), containing mostly floating (plankton) or swimming (nekton) organisms;

- the littoral zone (shallow water with light penetration), mainly containing rooted plants and nekton, also benthos.

In marine environments a comparable classification can be made, adding the tidal zone near the coast.


Horizontally also patterns are recognizable. Just like the dispersion in populations, in communities different population dispersions occur, namely (1) random, (2) regular (uniform), (3) clumped (contagious), and (4) regular distribution of clusters.

Again the contagious dispersions (3) and (4) are most common. The horizontal pattern probably results from various factors. First, the physi

cal environment does not allow settlement on all sites; particularly light intensity, moisture and soil structure may be limiting. Second, dispersal from parents may lead to clusters of young individuals (this holds both for plants and for animals). Third, species may be interrelated, as described in earlier sections.

The structure of a community is also influenced by temporal factors. In forests or coral reefs diurnal and nocturnal animals are active at different times of the day. In aquatic communities planktonic animals migrate up and down in daily cycles. In different seasons, different communities prevail in both aquatic and terrestrial environments. For instance, in fresh water diatoms are dominant in winter, desmids in early summer and blue-green algae in high summer.

In conclusion, in a community many species find their place as a result of spatial and temporal opportunities. The community apparently provides a number of niches, the amount of which is determined by the variety of environmental factors governing the community.

Species Diversity and Dominance

In a community there may be many or few species present. This species richness is usually called alpha-diversity. In general, species richness decreases from lowlands in tropicallatitudes to the polar regions and to the highlands. In the arctic regions, only few species are able to cope with the rigorous climate; these species have therefore relatively few competition problems. In a tropical rainforest or a coral reef environmental conditions are more stable, enabling many species to survive there. However, they are subject to interaction with other species.

Of course there are big differences in alpha-diversity within one region. Forests often have a high species richness because of their vertical stratification. Grasslands which are moderately grazed are rich as well. Different groups of species show different diversity trends: for birds and insects vegetation structure is important, for plants temperature and moisture are decisive, benthos is favored by relatively stable environmental conditions.

Apart from the number of species, the numbers per species are an important feature. This is usually called the evenness or equitability component of diversity. For instance, there is a great difference if in a community all species have equal numbers, compared with the situation that only few species have high numbers and the rest is rare.

The differences between common and rare species are dealt with by the concept of ecological dominance. In a community there are often only a few species which are really abundant (in numbers, biomass or productivity). These species may be called the "key industries" of the community because they channel most of the en-


ergy and nutrient flows. The reason for such dominances is not always c1ear. In newly formed habitats, good "colonists" (species with a strong r se1ection) may be dominant initially, but in due time they are repressed by K selection species. In other cases the question who comes first may be determinative for dominancy.

Common diversity indices combine both the variety and the equitability component:

c = L: (~r (Simpson's index)

H= - L (~)IOg (~) (Shannon-Wiener index),

in which n i = "importance value" for each species (numbers, biomass, production etc.)

N = the total of importance values. Simpson's index gives most weight to common species; the Shannon-Wien er

index stresses the importance of rare species. In the countryside of the temperate regions (parts of Europe and Northern

America) human influence used to enhance diversity of plants and animals. Modern agricultural technology has led to uniformity of arable land and pastures, resulting in low diversities.

Communities Along Environmental Gradients

The presence of a population of a species at a certain time in a certain place is the result of the interaction with other species and with environmental factors. This, however, does not necessarily inc1ude that distinct, c1early defined communities can be discerned. The so-called individualistic hypo thesis states that species are individualistically distributed and do not form well-defined communities. Thus any observed community would be merely a result of the coexistence of species. The opposite viewpoint inc1udes that species are co-evoluted, thus being more successful if in assemblage than without any obvious relationship.

The general pattern of community change along an environmental gradient is the replacement by each other of dominant species. The accompanying species (i.e. the majority) composition also changes along the gradient but this is less conspicuous. Apart from the alpha-diversity (within the habitat of a community), beta-diversity can be observed, being the diversity "between habitats" along a specific environmental gradient. Acutally there are three types of gradients which can be analyzed [57]: (1) the environmental gradient (also called complex-gradient) which results from

abiotic factors plus the influences of organisms on the abiotic environment; (2) the community gradient or coenocline; (3) the gradients of communities and environments together (i.e. of ecosystems),

called the ecocline. Major ecoc1ines are the temperature gradients from the tropics northward in

forests, c1imatic moisture gradients from tropical rainforests to deserts and the rocky ocean shores. Major communities (biomes) will be reviewed in the last Section.


The specificity of some communities probably has a partially evolutionary character. In due time species may change their habitat significantly; also particular species interrelationships may develop. This could explain why not all communities are "perfectly fitting perfect environmental gradients".

Ecosystems

Introduction

An ecosystem is the assemblage of organisms and their abiotic environment together. Communities consist of organisms only; the energy used to power the community system and the circulation of matter form extra dimensions. An ecosystem is a functional unit; the components are connected by energy and matter flows.

It is convenient to use the word ecosystem only in a functional way, in order to understand the processes which connect the components. The spatial-temporal form of an ecosystem is often called an ecotope. Thus, if we see a pond we may call it an ecotope; if we study structure and functions of the ecotope, we should speak

Soi I

8 // \@<:!> // \ ...... 0 " \ &

\

\ \ I B

\ I '.':' .: '~';..' .... . ~

.1 ,." (; .... ,' .. .' ....

~ ---------------- -~ )I -:::::-

-A----. \

~C2 ...... Sedi m e n ts I

~

----------- - \ ). --- - --- - - - -----

= Parent material-- - \ I I

===Parent material \ /

o@ Fig.9. Structure of a terrestrial and an aquatic ecosystem (After Odum [38]). A Abiotic substances. B Producers (vegetation on land, phytoplankton in water). C Consumers in three groups: Cl grazing animals (mice on land, zooplankton in water); C2 saprovores (detritus-feeding invertebrates); C3 carnivores (birds and fish). D Decomposing bacteria and fungi


of the pond as an ecosystem. The major components of ecosystems have already been discussed: on one hand, the organisms, divided into autotrophs, heterotrophs and saprotrophs, on the other hand the abiotic environment, consisting of inorganic substances, organic compounds and the climate regime. The system as a whole is steered by input of any form of energy. In Fig.9 the structure of a terrestrial and an aquatic ecosystem is shown.

Ecosystems may change in time; they may develop, get "mature," they may remain in a stationary state or decline. This change in time will be discussed in the next Section.

Trophic Structure

Energy is transferred through the ecosystem through pathways formed by organisms. The repetition of "eating and being eaten" [37] is called a food chain. There are two types offood chains: (a) the heterotrophfood chain which runs from green plants via herbivores to carnivores; (b) the saprotroph food chain leading from organic detritus via microorganisms to detritivores and their predators. A food chain is a linear sequence of species, such as

plant --+rabbit --+ fox. Another type of heterotroph food chain is the parasite chain, e.g. bird --+ flea --+ protozoan parasite. If more species are using a variety of resources, this "who eats whom?" is called

afood web. The more species present in an ecosystem the more complicated such a food web will be. However, as regards energy flow food webs are often characterized by a few species being dominant: the "key industries."

To get an overall picture of the trophic structure of ecosystems, it is often convenient to place all species in trophic levels, being groups of species which are an equal amount of "food steps" away from the primary producers (the green plants):

Primary Producers Primary Consumers

(= Secondary Producers) Secondary Consumers Tertiary Consumers

Green plants Herbivores

(Primary) carnivores, Parasites (Secondary) carnivores, Hyperparasites

Trophic level 1 Trophic level 2

Trophic level 3 Trophic level 4

Food chains of more than four steps rarely occur; a fifth trophic level does not really exist, also because many omnivores (feeding on both animals and plants) feed on two or even three levels. The same holds for the saprotroph food chain.

Trophic levels are not intended to classify species; they are only useful in understanding the flows through the system. By means of consumption, materials are constantly circulated (except for the input-output relationship with the surrounding ecosystems). Energy is channeled through the trophic levels; much ofit is used for "maintenance work" like respiration (often up to 90%). Thus in a food chain or through trophic levels there is a considerable loss of energy. This explains at least to a great deal the form of the Eltonian pyramids (Elton [12]; originally they were called pyramids ofnumbers), ofwhich an example is given in Fig.lO. As this Figure shows, besides numbers various units for constructing an ecological pyramid are possible, e.g. energy flow, biomass or productivity. E. P. Odum [37] sug-


(d) .1

(c) 1.6

(b) I 14.1 I I 21.2 (E!)

(a) I 87.1

Fig.lO. Eltonian pyramid of energy flow (MJ (m2(year) in Silver Springs (After Odum [39]). a Primary producers, b Primary consumers, c Secondary consumers, d Tertiary consumers, e Saprotrophs

gests that numbers overemphasize the role of small organisms while biomass does the same for large organisms; energy flows (i.e. production + respiration) seem most suitable for comparing components of an ecosystem.

Production

Primary productivity of an ecosystem can be defined as the rate at which radiant energy is stored by photosynthetic or chemosynthetic activity of autotrophs. Gross primary productivity inc1udes the fixed energy used for respiration during the measurement period; net primary productivity is the rate of "definite" storage. Biomass is the result of net productivity. In Table 3 some figures for primary production of major world ecosystems are shown.

In terrestrial ecosystems there is a tendency for productivity and for biomass to increase along moisture gradients. Deserts and tundras have biomasses of 0-2 kg/m2 , grasslands 0.5-3, shrublands 2-10, woodlands 4-20 and full-grown forests 20-60. Measurements of biomass are usually done by sampling fractions (sterns, branches, leaves, bark etc.) and determining dry weights. Assessment of gross productivity is done by measuring the CO2 exchange (inc1uding both uptake and respiration) of plants.

Secondary productivity is the rate of energy storage at consumer (heterotroph and saprotroph) levels. Usually the terms net and gross are not used, secondary productivity being comparable to net productivity of plants.

Table 3. Global net primary production per year of some major ecosystem types. (After Whittaker [58], modified)

Ecosystem

Open ocean Reefs, estuaries etc. Desert and semi-desert Savannah Temperate and boreal forest Tropical forests Cultivated land

Surface area (106 km 2)

332 2

42 15 24 24.5 14

Global net primary production (10 12 dry kg(year)

41.5 3.7 1.7

13.5 24.5 49.4

9.1


Fig.ll. Pathways of food use by animals

In animals, only a small proportion of the energy and nutrients ingested leads to increase of biomass. Figure 11 shows how food and energy are used. Therefore the contribution of animals to the biomass (or standing crop) of an ecosystem is low in comparison with the autotroph biomass. Moreover, animals only use about 10% of the net primary production. The remaining 90% is, sooner or later, converted into inorganic and organic substances by saprotrophs. This decomposition along the detritus food chains in an ecosystem is necessary for the recycling of nutrients to the autotrophs. In terrestrial ecosystems, saprotroph biomass may be fifteen times that of animals. In marine environments the detritivore biomass may even be twice as much as the autotrophs biomass.

With each step along a food chain losses of energy occur. Ratios between energy flow at different points along the food chain, expressed as percentages, are called ecological efJiciencies. Within a trophic level, the growth efJiciency can be expressed as

net productivity at level i energy intake at level i .

Between trophic levels, the consumption efficiency can be written as

energy intake at level i net productivity at level i-I .

Many other efficiencies for communities have been proposed, using also assimilation rates, absorbed light, total photosynthesis, respiration. A review is given by Kozlovsky [21]. However, in practice it is very difficult to measure efficiencies of communities. Much more is known about food chains consisting of ecological dominants.

Biogeochemical Cycles

The importance of chemical substances for organisms was already indicated in Sect. 2. These substances circulate in the biosphere via pathways, called biogeochemical cycles, from environments to organisms and back. This circulation may be local (within an ecosystem) or global. Substances are transferred from one


- energy flow

nutrient - -+ cycling

~ energy loss -:- (respiration etc.l

I

I

I \

Fig.12. Energy flow and nutrient cyc1ing in a hypothetical ecosystem

27

ecosystem component to another; there is a continuous exchange of material between organisms or between organisms and the abiotic environment.

The flux rate is defined as the quantity of a substance transferred per unit time and per unit volume of the system. However, not all substance is continuously moving: most of it is in relative rest within a biogeochemical pool, being an ecosystem component. Such a pool can be the sediments at the bottom of a lake or the standing crop of green plants of a forest. The latter situation is the case in tropical forests. If these are cut, the biogeochemical pools simply disappear, in contrast with temperate forests where the soil forms the main nutrient pool.

An illustration of ecosystem energy flow and biogeochemical cycling is given in Fig.12.

Two concepts are useful in assessing the importance of flux rates in relation to pools. First, the turnover rate:

flux rate (in or out) nutrient quantity in pool·

Second, the turnover time, being the reciprocal of the turnover rate, indicating the time required for movement of a nutrient quantity equal to that in the pool.

There is considerable difference in turnover times in the various components of ecosystems. Generally, biotic components (particularly heterotrophs) have short turnover times, abiotic components (like sediments) often show very long turnover times. Biogeochemical cycling on aglobaI scale (including weathering ofrocks etc.) is therefore to be measured on a geological timescale sometimes.

Two major types of biogeochemical cycles can be distinguished. First the gaseous type in which the largest reservoir is the atmosphere or the hydrosphere. Examples are the carbon, nitrogen, oxygen and hydrological cycle. Second, the sedimentary type in which the reservoir is the earth's crust. Examples are the sulphur, phosphorus and iron cycle. Most cycles of the micronutrients are of the sedimentary type.

It may be clear that in most cycles the geochemical process is dominant rather than the biological regulation; most of the important reservoirs are abiotic. The same holds for the driving forces: in the hydrological cycle the solar energy (evap-


oration), in the phosphorus cycle oceanic upwelling. However, particularly the autotrophs and many species ofbacteria perform important roles in the major biogeochemical cycles (e.g. nitrogen, sulfur, phosphorus). The functioning of the biotic components of the biosphere strongly depends on their possibilities to control the biogeochemical cycles.

Influences like pollution may disrupt the turnover balances, thus changing nutrient supply. Natural systems, however resilient, may not be able to resist such a change.

Succession and Steady State

Succession

All ecosystems change in time. Lakes are filled up with silt, become a marsh and then a grass land or even a forest. Bare sand dunes are initially invaded by small resistent plants, then develop a turf which may eventually change into shrubland or woodland. This process of changing in vegetation (or, broader: community) structure is called succession.

Causes for succession can be twofold: autogenie, if the components of an ecosystem change their own environment; or allogenic, if forces from outside change the system. In most successions both causes are involved. Anyway there is a gradient of ecosystems in time: "a succession is an ecocline in time" (Whittaker, 1975, p. 171 [58]). If succession occurs in an environment where previously another community has been established (as is the ca se in areas devastated by fire), this is called a secondary succession. In newly-formed habitats (starting with a bare soil) we speak of a primary succession. In Europe, only a few primary forests (dating from the end ofthe last glacial period) exist: most forests are secondary, ifnot predominantly artificial.

The different communities following after each other form asere. However, ecosystems do not always develop linearly: sometimes cyc/ic succession occurs. In such a cycle four main stages can be observed: pioneering, building, maturating, and degenerating stages. Watt [53], in a study on Calluna heath, observed that the dominant plant, Calluna, is degenerating by age. It is replaced by lichens which, after dying, leave the soil bare. Then Arctostaphylos invades the area which is then overpowered by newly invading Calluna. The whole cycle takes about 50 to 60 years.

Together with the development of the community the soil properties in terrestric environments change. Depth and organic content increase and different layers develop. Nutrient pools in the soil as well as in the vegetation increase. The aboveground development of vegetation causes a change in microclimate which enables other species to settle. The rate of change in species composition slows down as the community is developing.

The Climax Concept

The final stage of a succession is called the c/imax. In that stage, the community is at maximum stability under the present environmental conditions. Tropical rain


forests are examples of climax eeosystems, but bare yellow dunes are as weIl: in both examples environmental eonditions are exploited maximaIly.

Clements [3] first produeed the eoneept that has beeome familiar to eeologists as the "monoclimax." In his view, climate is the only environmental eondition whieh really governs sueeession. Any disturbing faetor (topography, soils, drainage) is thought to change in due time and not to override the effeets ofthe climate. Different climax vegetation eompositions due to disturbanee ete. were ealled disclimax or subclimax.

Other eeologists have developed the "polyclimax" hypothesis. Tansley [51] and Daubenmire [8] stated that in a given area different climaxes may be found, eontrolled by soil moisture, nutrients, topographie faetors, animals or wildfires. The last faetor ean also be regarded to eause a "eyclie eatastrophie climax" by removing most of the biomass, as is the ease in Californian eoastal ehaparral eommunities.

FinaIly, a "climax pattern" hypothesis was launehed by Whittaker [56]. This hypothesis is based on the eommunity eontinuum along environmental gradients and allows for a eontinuity of climax types varying along a gradient.

Table 4. Trends in the development of ecosystems. (After Odum [36])

Ecosystem attributes

Communityenergetics

1. Gross production/community respiration (P/R ratio)

2. Gross production/standing crop biomass (P/B ratio)

3. Biomass supported/unit energy flow (B/E ratio)

4. F ood chains

Community structure

5. Species diversity - variety component 6. Species diversity - equitability component 7. Biochemical diversity 8. Stratification and spatial heterogeneity

(pattern diversity)

Life history

9. Niche specialization 10. Life cyc1es

Nutrient cycling

11. Mineral cyc1es 12. Nutrient exchange rate, between organisms

and environment 13. Role of detritus in nutrient regeneration

Selection pressure

14. Growth form

15. Production

Developmental stages

Greater or 1ess than 1

High

Low

Linear, predominantly grazing

Low Low Low Poorly organized

Broad Short, simple

Open Rapid

Unimportant

For rapid growth ("r -selection")

Quantity

Mature stage~

Approaches 1

Low

High

Weblike, predominantly detritus

High High High Well-organized

Narrow Long, complex

Closed Slow

Important

For feedback control ("K-selection")

Quality


Agricultural systems which have been managed in the same way for a long time may be called anthropogenie (sub )climaxes. Of course this only holds if the system is stable; modem fossil-fuel subsidized agriculture does not produce any coherent system of plants and animals at all.

"Y oung" ecosystems which rapidly develop have other characteristics than "mature" ecosystems which change only slowly or not at all. Odum [36] has produced interesting information about these differences. Parts of it are shown in Table 4. We shall briefly discuss some of these topics.

As regards production, the PjR ratio of developing communities often markedly differs from 1. In autotrophic succession PjR is larger than I; this situation holds for developing grasslands etc. On the other hand, in polluted streams with a heterotrophic succession PjR ratio's are smaller than I, whieh means that respiration exceeds increase ofbiomass. PjR ratio's in stable, developed systems equal 1 approximately; these ratio's therefore are a good indication of "mature" or climax systems. It also means that mature systems use most of the energy flow for maintenance and not for increase in biomass.

All kinds of diversities are high in mature systems, at least compared to earlier developmental stages. Weblike food chains presuppose a great variety of species within the different trophic levels; most of these species are not dominant (high equitability). Vegetation in mature systems is very stratified, allowing for many niches for animals. Nutrients mainly circulate within the system; the exchange with the surrounding ecosystems is relatively unimportant. The detritus food chain is very complicated, producing most of the nutrients for the standing crop of vegetation.

Species in mature systems tend to have long, complex life cycles occupying many "narrow" niehes. They mainly belong to the group of "K selection" species: quality is more important than quantity.

Steady State and Stability in Ecological Systems

In ecological theory the concept of the "steady state" of systems is frequently used as regards structure and function ofecosystems. H. T. Odum defines a steady state as a situation in which a constant pattern of flows, cycles, storages and structures prevails. Inputs must equal outputs, particularly of energy [40].

Homeostasis of ecological systems is the tendency to resist changes and to remain in an equilibrium state [37]. This concept also regards a steady state. Derived from cybernetics, it assumes feedback controls to be at work in ecosystems. Positive feedback causes the system to accelerate because processes are stimulated by themselves. Rapid population growth of species in newly colonized habitats is a good example. Negative feedbacks slow down processes thus counteracting deviations, as is shown by predator populations which increase when the prey population grows.

Theoretieally, steady state situations may occur at all ecological organization levels because of the fact that usually many (biotic and abiotic) factors are involved. However, particularly climax ecosystems often show a remarkable constancy over long periods. During succession ecosystems develop into a certain direction; in climax situations changes are fluctuations around a certain level.


The concept of the c1imax is not identical to the concept of the steady state. The c1imax is a final stage of a succession; nevertheless the c1imax community may be very unstable because of rigorous environmental factors (e.g. in tundras or deserts). Such a system can possibly be called "mature" but is not homeostatic as the system strongly depends on exogenous influences. On the other hand, tropical rainforests and coral reefs are examples of steady state systems which are self-sustaining.

Ecosystems may be in steady state for many years, but probably not infinitely. At first, the cyc1ic wildfire succession may inc1ude a stable c1imax for 20-200 years which is nevertheless ended time and again. Second, other "catastrophes" like volcanic eruptions or earthquakes may destroy c1imax ecosystems, causing secondary succession. Third, ecosystems change with age, particularly after more than 200 years; dominance relationships then change, causing a decrease in primary production. Fourth, c1imatic changes, particularly from humid to dry, may change forests into woodlands, grasslands or even deserts. Fifth, evolutionary changes ofthe species may change the ecosystem. Sixth, human influences may have decisive effects on the steady state (see the final Section).

Although an ecosystem may be considered to be in a stationary state over a long period and on a large surface area, many short-term fluctuations on a smaller scale may occur. Odum reports [39] that the Silver Springs ecosystem is in a steady state on an annual base. However, primary production is markedly re1ated to seasonal changes, which also holds for the population sizes of some of the animal species. The same goes for most temperate ecosystems (forests, grasslands).

All in all, the relative stability ofmany ecosystems, although measurable, is difficult to explain. Particularly in the sixties ecologists have hallowed the diversitystability-hypothesis which basically states that increasing diversity causes increasing stability. This hypothesis appears to be supported by the stability oftwo ofthe most species-rich ecosystem types ofthe world: tropical rainforests and coral reefs. In the last decade, however, many ecologists have questioned this hypothesis. Stability should be rigidly defined as the concept may have different meanings like constancy (= resistance), persistence, inertia, e1asticity (= resilience), amplitude of regular fluctuation and so on. Many types of system stability may occur together, thus rendering difficulties in deciding what causes what. Diversity or complexity, on the other hand, are re1ative1y simple concepts, indicating the number of functional components within a system.

Probably the question what causes (in)stability is not quite correct, if asked from the ecosystem point ofview. Abiotic environmental conditions (like c1imate) may allow for fragile (few perturbations possible) or only for robust (resistant to severe perturbations) systems. Thus boreal forests (low diversity) may stand both a harsh c1imate and other important influences, while tropical forest systems (high diversity) may not be able to counteract even slight perturbations.

In this context the evolutionary changes within systems become extremely important. Species may become adapted to changed environmental conditions (sometimes caused by other organisms) via adaption ofpopulation growth strategies, of nutrient requirements and so on. Adaptation and evolution are continuous processes occurring everywhere in the ecosystem. In "permitting" environments species richness tends to increase for a very long period because of many small adjust-


ments of niches. Under rigorous environmental conditions such adjustments are not possible because all species need a "wide niche space" to survive.

Following May's conclusion [31] on behalf of experience with model ecosystems, a predictable ("stable") environment may permit a relatively complex and delicately balanced ecosystem to exist; an unpredictable ("unstable") environment is more likely to demand a structurally simple, robust ecosystem. Further discussions of these concepts can be found in Van Dobben and Lowe-McConnell ([10], pp. 139-236) and May ([31], pp. 142-162).

Major Ecosystems of the World

Introduction

In the foregoing Sections a variety of principles regarding structure and function of ecosystems has been discussed, using the organization levels as a starting point. However, most ofthis cannot be seen "in the field," as ecological processes require time to occur. In this Section attention will be given to some general structural features of major ecological systems. For convenience, terrestrial, freshwater and marine systems will be discussed separately. As terrestrial ecosystems are most conspicuous, they will get major attention.

The general structure of major ecosystems consists of different local ecosystems containing several trophic levels. Within these levels many populations of plants, animals and microorganisms occur. Although the species composition of ecosystems may vary from place to place, overall limits are set mainly by major abiotic conditions. In terrestrial environments, temperature and moisture are most decisive, in aquatic environments temperature, light and physical forces like tidal movements are most important.

Terrestrial Ecosystems

Structure

The assemblage of all plants in an area is called the vegetation. The form and composition of the vegetation is often very indicative for the abiotic "possibilities" of the area; therefore terrestrial communities can be indicated by using vegetation classification. Theflora is the assemblage oftaxonomic units in an area, containing vascular plants, mosses and lichens. In many communities the species list is rather long, while only few dominating plants establish the structure of the vegetation as a whole.

As regards animals, vertebrates and insects are presently most dominant in terrestrial ecosystems. Animals may vary from sessile species to extremely mobile ones. This of course weakens the concept of the ecosystem; particularly migrating birds belong to more than one ecosystem. However, just as the system sustains plant species which grow and flower only in early spring or in late summer, it also sustains animal species which are not present in all seasons. The same holds for animals only active during the day or during the night. Apart from temperature and


moisture, the soil is probably the most important abiotic factor for terrestrial systems, being the source of most nutrients. In itself, the soil forms aseparate ecologieal subsystem where decomposers convert dead biomass into available nutrients anew.

The group of the decomposers consists of fungi, (heterotrophie) bacteria, actinomycetes (fungi-like bacteria) and protozoa. Fungi and bacteria decompose relatively simple organic substances like aminoacids and sugar. Actinomycetes contribute to the forming of humus in the top soil. The breakdown (or mineralization) ofhumus is not yet quite understood; both microbes and abiotie factors playa role.

Decomposers plus soil algae (autotrophs) together form the microbiota of the soil. A group oflarger animals (mesobiota) includes nematodes, mites, springtails (all very abundant), small enchytraeid worms and insect larvae. This group mainly feeds on bacteria, algae and plant roots. The group of macrobiota consists of plant roots, earth worms, larger insects, spiders, snails and vertebrates like moles. This group is partly autotroph, partly herbivorous and partly carnivorous.

Biomes

The major terrestrial ecological formations or biomes are defined in accordance with regional climates. Within a biome the form of the climax vegetation is supposed to be uniform. Of course many successional stages of the climax may occur within the biome; grassland communities, for example, may be developmental stages within the deciduous forest biome.

Biomes have also distinct soils and faunas, although overlap frequently occurs. The following description of some biome groups therefore also partly deals with these aspects. In Fig. 13 the biome groups and soil types are arranged according to temperature and moisture conditions.

1. Tropical Forests. This group includes rain forests (rainfall weIl distributed throughout the year) and seasonal forests (in humid tropical climates with a dry season during which many trees lose their leaves). Trees are tall, ofnumerous speeies; The canopy is also rieh of other plant and animal species, notably orchids, mammals, reptiles and invertebrates.

One of the major soil types in these biomes is the red forest soil (rain forest latosol). Such soils have been intensely weathered and leached out, resulting in nutrient-poor substrates for plants. Therefore tropical forests have a nutrient circulation whieh is almost independent of the soil. If the forest is cut, the clay of the soil (containing ir on oxides) hardens if exposed to air, thus forming the so-called laterite.

Tropieal forests occur in South and Central America, Central and West Africa, and South East Asia up to North Australia.

2. Temperate Forests. This group contains rain forests (the giant coniferous forests in California), deciduous and evergreen forests (on all continents though mainly in East North America, Europe and East Asia). Climates vary from cool maritime (much rainfall) in the rainforests to summer-dry in evergreen forests.

34

-15 -, r 6 1 u ,",

o

ai -5 ':> -:;; 0 'QJ

~ 5 ~

r ..... , 4 \

I' \ ",---IV I \ \ , , ~ 1

'\(\\ \ 2.b I

co 15 1 7 1 \ 5 \ \ - - - - .-IV I I \ \ d \ \ :;; I e I \ \ ~ -- ---QJ \ I \ - \

:E: I I 25 \~ I \ 4

,/

/

I

\ 1. a

..... - - --100 200 300

Mean annuat precipi tation, cm

, I

/

400


Fig.13. Major biome types and soils in relation to mean annual precipitation and mean annual temperature. Legends to numbers: Biomes: 1 tropical forests, 2 temperate forests, 3 taigas, 4 woodlands and shrublands, 5 grasslands, 6 tundras and arctic-alpine systems, 7 deserts. Soils: arainforest latosol, b forest brown earth, c taiga podzol, d chernozem, e sierozem

Species diversity in temperate forests is much lower than in tropical forests. Tree heights vary from over 100 m (Californian redwoods) to 20 m (European deciduous forests).

In many temperate forests the so-calledforest brown earth occurs. This soil type has a thin litter layer (mull), is weakly acid and contains much silica. Brown earth is relatively nutrient-rich; the nutrient pools ofthe temperate forest ecosystems are mainly located in the soil, not in the plants.

3. Taigas. These needle-Ieaved forests are on the cold "climatic edge" offorests in North America and Eurasia. Species diversity is low; forests may change into woodlands or shrub communities near the northern tree line. Animals like moose, lynx, nuthatch and warbier species are dominant.

Taiga podzols form a major soil type. Decomposition is slow, the rather thick litter layer is called mor. Beneath the mor a bleak leached layer occurs (hence podzol- Russian for ash), under which a hardpan of iron and humus is found. This soil type is nutrient-poor, the main nutrient stocks being present in plant biomass and litter.

4. Woodlands. In these biome types trees are not forming vast forests but are scattered. Tropical broadleaf woodlands (mainly in Central Brazil and Central Southern Africa) and temperate woodlands (Western United States, Mediterranean, Southern Australia) occur in climates too dry for true forests. Dominant trees may be needle-Ieaved, sclerophyll or deciduous broadleaved. In still more acid climates woodlands give way to the thornwood (Central South America) and temperate shrubland (Mediterranean maquis, Califomian chaparral and Australian mallee), biomes which are often in a cyclic wildfire succession.


5. Grasslands. Subtropical grasslands or savannahs mainly occur in Australia and in Africa where they support an enormous diversity of grazing animals. Temperate grasslands occur in vast areas of Eurasia (steppes) and North America (prairies) where a continental climate prevails, but also in South America (pampas), Africa (veldt) and Australia. Plant species diversity is rather high, compared to temperate forests, but animal species richness is much lower because of the lack of different vegetation layers. Mammal fauna includes small rodents and large herbivores like bisons, zebras, antelopes and kangaroos. Grasslands are also frequently subject to wildfire.

Woodlands and grasslands often stand on chernozem soil (black earth). This soil type is not very much leached because water use and evaporation often exceed water influx. It is thus fertile; large amounts of organic compounds (from dead plant roots and animals) are also present. The sustained biomass is not very large but has a rapid turnover; most nutrient pools are in the soil itself rather than in the vegetation.

6. Tundras and Alpine Biomes. Tundras (North America, Eurasia) are treeless arctic plains dominated by small shrubs, sedges, grasses, mosses, and lichens. The fauna includes reindeer (caribou), lemming, arctic hare, plovers etc. The deeper layers of the soil are often constantly frozen, so that life only occurs in the topsoil and above ground.

Alpine shrublands and grasslands occur above the timberline. Shrub communities may be dominated by heaths (Africa, Himalaya) and by composites (Senecio; Africa, New Zealand). Most alpine biomes are grasslands, often dominated by sedges, grasses and cushion or rosette plants. Many animals active in winter are colored white (ptarmigan, mountain sheep).

7. Semi-deserts and Deserts. These biome types occur in arid climates. Warm semi-desert scrubs are found around the deserts in North America, Northern Africa, Southern Asia, and Central Australia. Succulent plants dominate (creosote bush, saguaro); there is a rich lizard and snake fauna. Cool semi-deserts occur in the western United States, Central Asia and Southern America. These ecosystems are dominated by plants like sagebrush (Artemisia). Arctic-alpine semi-deserts with mainly spiny cushion plants range from the Mediterranean into Central Asia, but are also found in arid regions of the tundra biome.

Deserts mainly occur in subtropical zones where rainfall is sometimes less than 2 cm per year. Vegetation is very spar se or lacking. Thus the physiognomy is determined by the ground surface: sand, rocks, stones or salto Deserts are found in Asia (Thar, Gobi), Western Australia, Northern America (Sonora, Mojave) and Africa (Sahara, Kalahari). Animallife is also very poor, in contrast with semi-desert areas.

Arctic-alpine deserts occur in the coldest climates where water is less limiting life than temperature. Only few plant and animal species can stand these conditions.

Sierozem, or grey desert soil, is a major soil type in (semi-) desert communities. There is almost no litter layer; nor are plants influencing the soil very much. Upwelling water is rapidly evaporated, transporting many nutrients to the soil surface. As the biomass sustained is very low, nutrients are mainly stocked in the soil.

36

Freshwater Ecosystems

General

Two features of freshwater ecosystems are most characteristic: (a) the substrate of the communities is predominantly water, (b) tbis water lacks large concentrations of salts.


As water has particular thermal properties the temperature range is narrower than on land. Nevertheless temperature is an important limiting factor as aquatic organisms are often stenothermal. Salinity in freshwater is always less than 0.5 promille. Important nutrients like nitrates, phosphates, and calcium are often limiting factors. Lack of salts in freshwater environments has led to specific adaptations of organisms in regard to osmoregulation.

Freshwater ecosystems are in a continuous succession because of sedimentation or erosion. These processes often can only be measured on a geological time scale. Nevertheless aquatic communities are constantly in adaptation because of slight changes of environmental conditions.

Freshwater ecosystems may be divided into two broad categories: 1. Lentic communities in which the water is not rapidly moving: lakes, ponds,

swamps or bogs. 2. Lotic communities in wbich running water prevails: rivers, streams or springs.

Lentic Communities

The zonation in ponds and lakes has already been shown in Fig. 8. Autotrophs in the littoral zone are mostly rooted seed plants and floating algae. The latter group may cause huge "blooms" in case of excess of nutrients through pollution. Diatoms, green algae and bluealgae are dominant groups. The rooted aquatics can be divided into emergent plants (photosynthesis above water surface), plants with floating leaves (assimilation at surface) and submergent plants (assimilation under the water surface). In the limnetic zone autotrophs are phytoplankton, consisting of the three groups mentioned plus dinoflagellates. Most species are microscopic and therefore inconspicuous. In the profundal zone no light for photosynthesis is available.

In the heterotroph group invertebrates are by far dominant, particularly arthropods. The littoral zone contains most species; these are largely different from those of the limnetic zone. Fish move freely between the zones but spend much of their lifetime in the littoral zone, e.g. for breeding. Amphibious vertebrates like frogs and water snakes alm ost exclusively occur in the littoral zone. In the limnetic zone copepods, cladocerans and rotifers are dominant. The profundal zone mainly consists of detritus feeders, bacteria and fungi wbich play an important role in "regenerating" nutrients.

In lakes the limnetic and profundal zone are relatively large, if compared with ponds. Lakes are clearly stratified as regards temperature and oxygen content. Heating of the surface water (epilimnion) in summer causes a sharp thermocline towards the lower layer (hypolimnion) where the cool water does not circulate; oxygen supply becomes depleted. In winter the hypolimnion becomes the warmest


stratum, causing circulation of water and renewed supply of oxygen. In warm climates less conspicuous turn-overs occur.

Lakes are often classified according to their depth and their nutrient contents. Generally oligotrophie lakes are deep, having a low primary production and few plankton blooms because of lack of nutrients. Species richness is often very high, particularly in ancient lakes like lake BaikaI. Contrarily, eutrophie lakes are shallower and have a high nutrient content and a large primary production. Because of sharp turn-overs, species with narrow tolerance ranges are easily extinct.

Lotie Communities

Running-water ecosystems show three major differences with lentic ones. First, currents are limiting the occurrence oflife-forms, particularly in the centre ofwater beds. Second, the substrate of river beds is very important as a fixing place for most organisms; nutrients are flotating by rather than encountered. Because many nutrients come from terrestrial or lentic ecosystems, lotic systems are heavily dependent on these. Finally, oxygen supply is always large; organisms in running water are very sensitive to reduced oxygen.

Lotic communities can be divided into rapid and pool communities. Organisms in pools are often the same as those in ponds. In rapids plankton also occurs, although this seems unlikely. It is supplied by ponds, lakes and pools in large quantities and travels from one section ofthe stream to another. Sessile organisms show adaptations to the running water like permanent attachment mechanisms, hooks, suckers, streamlined or flattened bodies and positive rheotaxis (i.e. the capability to orient itself upstream). Organisms are also adapted to part of a longitudinal zonation of environmental factors (temperature, nutrient supply).

Marine Ecosystems

Seas and oceans have important functions in relation to the world's climates and the major biogeochemical cycles. They cover about 70% of the earth's surface; some are very deep with li fe in all depths. All marine environments are linked but major barriers are formed by temperature and salinity differences. Ofthe 35%0 salt in sea water almost 80% is NaCI; other important salts are K, Ca, and Mg compounds. In the oceans salinity is very constant; it may vary near land, where sea water is mixed with fresh water. Many nutrients occur in low concentrations, thus forming limiting factors to organisms; average productivity is therefore low. This does not hold for places where upwelling occurs.

Wind and variations in temperature and salinity cause eurrents of which the Gulf Stream, the North Atlantic Drift and the Humboldt Current are wellknown. Waves and tides form another type of movement of marine water.

Combinations of environmental conditions as mentioned above enable many different ecosystems to exist. It is difficult to distinguish "biome" types in marine environments but, by and large, the same division as in freshwater ecosystems may be used (see Fig.8). The word pelagic is commonly used for all non-benthic communities ofthe open sea, including plankton and swimming animals. Littoral com-

38

Intertidal : ! :

0

-1 E

..>::

.c 2 L ....

e. QJ

o 3

4

Neritic Oceanic

\ p

[S \

Sea level

b

AP


Fig.14. Zonation of marine ecosystems in relation to water depth. Legends: Intertidal: M mangroves, E estuaries, S rocky shores and sandy beaches. Neritic and Oceanic: e coral reefs, p pelagic, b benthos. Bottom: L land, es continental shelf, AP abyssal plain

munities on the continental shelf are called neritic in marine environments; the profundal zone is called aphotic ("without light"), in contrast to the euphotic zone (neritic plus surface pelagic communities). Near the coast the intertidal zone (between high and low tide level) is considered aseparate seetion together with sealand ecoclines.

Plant and animal groups in the sea clearly differ from those in freshwater. Seed plants are almost absent, except for sea grasses near the co ast. In phytoplankton red, green and brown algae dominate together with bacteria, diatoms and green flagellates. Animallife lacks the insects, these being replaced by crustaceans. Other important animal groups are coelenterates (jellyfish), echinoderms (sea-urchins), annelid worms, sponges etc. and of course fishes and whales. Birds feed on marine ecosystems but breed on land.

The zonation of the different marine ecosystems is shown schematically in Fig.14.

1. Shores and the Intertidal Zone. This group includes rocky shores, sandy beaches and mangrove swamps. All systems are subject to the regular fluctuation of the sea water level. Particularly at the low tide level rich communities occur which are exposed to air only during short periods. The remaining part of the intertidal contains less species which are adapted to a longer exposure to air like mussels and barnacles. These animals occur mostlyon rocky shores, accompanied by various algae groups. On sandy beaches the so-called interstitial fauna occurs, living between the sand grains. Other animals include the "sand burrowers" (isopods, amphipods, polychaete worms). Sandy beaches are very unstable environments because of the tide and the continuous movement of the sand.

Mangroves are emergent treelike plants that can stand the salinity of the sea. The roots reduce tidal currents and effect deposition of mud and silt. Thus mangrove ecosystems may extend the coasts, particularly on shallow shores.

2. Estuaries. Coastal waters which are partly closed off from the sea are called estuaries. This group includes coastal bays, river mouths (deltas, fjords) and tidal


marshes. Often estuaries are transition zones (ecotones) between freshwater and sea-water. The water may be stratified, fresh water flowing over salt water, or mixed because of vigorous tidal movements.

Species composition in estuaries varies according to geomorphology and climate. Tropical rivermouth systems are very species-rich, arctic fjords have a low diversity. Many species are endemie, i.e. restricted to the estuarine ecosystem. Juvenile animals of many fish and vertebrates are also found here, estuaries thus being "nursery grounds."

Production in estuaries is often very high. This is due to the function as a "nutrient trap"; also photosynthesis occurs almost the whole year. Next, tidal movement supplies energy for maintenance "work" in the ecosystem.

3. Coral Reefs. These ecosystems form fringes around islands, barrier reefs or hollow circular atolls. The coral itself is a coelenterate animal living in colonial structure in a "skeleton" of calcium carbonate, together with dinoflagellatic and filamentous algae. Primary production is very high, thus delivering food for many herbivorous species. Part of these are zooplankton but the amount of this is not enough to meet the requirements of the coral; the symbiotic algae supply part of the required energy and nutrients.

Species richness in coral reefs is amazing, as in the variety in forms and colors of animals and plants. The ecosystem of a coral reef shows a P IR ratio of near I, indicating a metabolie climax. Nonetheless production on all trophic levels is high, indicating a very efficient cycling of nutrients.

4. Neritie Eeosystems. This group is defined as the assemblage of ecosystems between the intertidal zone and the edge of the continental shelf. The floating or swimming (pelagic) group of organisms is dominated by phytoplanktonic diatoms and dinoflagellates, zooplankton (copepods, crustaceans: "krill") and large consumers like fishes. Many zooplanktonic species spend part oftheir life (particularly the adult stage) as benthos on the bottom of the sea. Species composition varies according to water depth and temperature. Neritic benthos show zonation from shallow to deep water. As the continental shelf is not a gradual slope, this zonation is mosaic-like. Distant communities at the same depth often are dominated by animals ofthe same genus (e.g. clams ofthe genera Maeoma and Venus). In shallow areas large kelp beds and sea grass "meadows'" occur which are used for food and shelter by pelagic organisms.

5. Oeeanie Eeosystems. This group is found outside the continental shelf. Surface pelagic organisms are comparable to those of the neritic ecosystems. Phytoplankton is mainly microscopic. Primary production per cubic meter is not large; however, because of the vastness of the oceans it is considerable (see Table 3).

As there is a continuous sink of nutrients towards deeper oceanic regions, deep pelagic and benthic animallife is also rich. Deep-sea species are adapted to high pressure, low temperatures and complete darkness. They often show particular features like "lanterns", enormous mouths or "tripods". The oceanic bottom surface is "mud", mainly consisting of skeletons of organisms and clay of volcanic origin. Most of the important invertebrate animal groups are represented by several specialized species.

40

Ecology and Environmental Problems

Introduction


"Environmental problems" are usually considered to be the non-desired changes in the environment of human beings. These changes are (also by definition) caused by human activities. Four types of "activities" causing environmental problems may be distinguished: 1. Addition of matter and energy to the environment, notably chemical substances

and radiation. This is called pollution. 2. Extraction of matter and energy from the environment, notably fossil fuels,

ores, timber. This can be called exploitation. 3. Destruction of structure and disturbance of processes in the environment, e.g.

building of dams or introduetion of non-endemie organisms. This can be called environmental disruption or perturbation.

4. Addition of more human beings to the world, i.e. population growth, which leads to an increase of activity per unit of space and time. Human beings have always used their environment. However, the core ofmany

environmental problems is mostly the intensity or rate of use of the environment, together with a rapid population growth. Environmental scienee studies the effects of human presence and human activities on the environment. Environmental changes mayaiso affeet the human population itself, e.g. in the case of pollution. The study of sueh problems is often ealled environmental health research. Environmental science is multidiseiplinary because environmental problems are exerting influence in many ways. During the last deeade specialisms like environmental chemistry and environmental economics have gained mueh importance. Ecology, being an environmental science in itself, has also been developed towards studying the effects of human society on the environment, notably nature.

Some environmental problems which have recently received attention of ecologists will be discussed briefly below. This discussion is meant to give a review of current issues rather than dealing with them thoroughly.

Pollution

Pollution can be roughly divided into chemical, thermal, radio-active and noise pollution. Emission of substanees or energy is mostly a local aetivity but streams, eurrents, and winds may spread a pollutant worldwide.

At the organismal and speeies level of organization pollutants may be simply toxic (many organic compounds), affecting the physiology ofthe organisms. Generally pollution can be seen as exceeding the toleranee range of an organism or speeies. If a particular nutrient (e.g. phosphate) forms a limiting factor already, discharge of it may eause excessive growth like algae blooms.

Some species are extremely sensitive to air pollution. Species oflichens, for example, disappear in regions where the prevailing winds carry pollutants. They can therefore be used as indieators.

Adaptation to relatively "sudden" discharges is not very likely, as evolution is a slow process. Change of composition of environmental factors means changes in


the niehe pattern; therefore different species, often with wider tolerance ranges, will take over.

Population growth is affected by pollution. Natality may decrease, mortality may increase and dispersal may occur as areaction on deteriorated conditions; r selection species have best chances. Predator-prey relationships and competition change because population densities change. Species with large tolerance ranges will prevail in competitive situations.

In terrestrial communities niches may disappear because habitat-supplying species cannot survive. Species richness decreases, dominants give way to other species; this also holds for aquatic communities. Ecosystems are changed because food chains and food webs are disrupted. Concentrations of many pollutants tend to increase along certain food chains ("biological magnification"); therefore large carnivores like birds of prey may consume lethaI doses of chemical substances. Biogeochemical cycles are disturbed as a result of excess of nutrients which cannot be managed. Nutrient pools tend to swell, thus buffering excessive quantities.

Succession may be halted by pollution. In extreme cases ecosystems may be "simplified" (less biomass, productivity, structural complexity, diversity); succession then starts anew, often in a completely different direction. Steady state systems give way to unstable systems tolerant towards a largely "unpredietable" environment.

All major ecosystems of the world are being affected by pollution. Terrestrial systems are mainly under the stress of air pollution, particularly sulphur and nitrogen compounds. Leaching of pollutants through the soil to groundwater mayaIso have effects. Aquatie systems have been most conspieuously changed by discharge of pollutants; "culturally" eutrophicated lakes and "dead" rivers are well-known examples. In marine environments the catastrophes with carriers of crude oil, effecting complete "ecocide" in the intertidal and neritic zones, have drawn most attention. However, sewage discharge into the sea still almost freely continues. The effects of using open water as garbage can are yet not at all understood but may, in due time, be disastrous for the present ecosystems.

Exploitation

Extraction of abiotic resources like fossil fuels and minerals as such has few ecological effects. Infrastructure and buildings needed for exploitation and use of the extracted resources cause environmental disruption and pollution.

Excessive exploitation of biotic resources may be disastrous at all ecological levels of organization. Species may disappear because scattered individuals cannot ensure continuous reproduction. Good examples are whale species and other large mammals, but also tropieal plant species. Excessive fishing has decimated populations of herring, sole and sardine; the use of nets with fine meshes leads to decrease of younger age classes.

Forestry in tropieal rainforest areas is threatening these ecosystems severely, particularly because the red forest soils may harden into laterite which is infertile. Erosion caused by overexploitation of forest has been a common feature in man's history (e.g. in the Mediterranean region). Modern forestry, using monocultures, is confronted with uncontrollable wildfires and outbreaks of pests. Moderate graz-


ing may be beneficial to species diversity in grasslands. Overgrazing in rangelands causes monotonous vegetations which are species-poor. In extreme cases overgrazing causes erosion and increase of desert areas, as is the case in the Sahel countries. Intensive farming in temperate regions requires use offertilizers which destroy nutrient balances and therefore change food webs.

The need for food for a growing human population has led to intensification of agriculture, e.g. by regulation of groundwater tables, leading to changes in flora and fauna. Monocultures are heavily subsidized by fertilizers and fossil fuels, with an increased chance on plant diseases and pests. New, high yielding varieties of crop plants are often much more vulnerable to diseases.

In general, exploitation of ecosystems is only possible if nutrient balances and food webs are not really disturbed. Overexploitation reduces and simplifies structure and processes within ecosystems, causing shifts in succession and sometimes definite destruction of ecosystems.

Environmental Disruption

Most of the effects mentioned above mayaiso be regarded as environmental disruption (an environmental problem should be seen as the complex of effects caused by a specific human activity). There are, however, other important influences which cannot be characterized as pollution or exploitation.

Waterworks influence major abiotie constituents of terrestrial and freshwater ecosystems, viz. water and soil. Barrages create new lakes on places where terrestrial systems occurred. Downstream the water course can be much more regulated. However, such regulations evoke other problems like outbreak ofpests (e.g. bilharzia in the Nile region after the Aswan dam had been built). In the "artificial" lakes plants like Eichhornia and Salvinia may bloom, just as algae. Oxygen content of the water may thus be strongly reduced, meaning death to other organisms. Other waterworks influencing ecosystems include canalization, drainage and reclamation ofland. In all cases the original ecosystems are destroyed, giving way to completely different (and often less stable) ones.

Urbanization and infrastructure activities have similar effects. Residential or industrial complexes convert natural ecosystems into artificial ones which have to be controlled at the cost ofmuch labor, fossil fuels, fertilizers and pesticides. Only very rarely are building activities compatible with ecosystem structure and function. Pipelines, high-tension cables and roads form artificial boundaries to ecosystems, causing isolation of individual organisms if not immediate death.

Finally, introduction of non-endemie species may threaten the existing flora and fauna. Introduced species often show easy adaptation to the new habitat; if present predators cannot control population growth of the "invader", population densities may cause different competitive relationships. Such situations are almost always detrimental to the native species.

Human Population Growth

Human population has increased enormously in the twentieth century. Improved medical care, labor conditions, food distribution etc. have contributed to this phe-


nomenon. The still growing population needs space, food, materials for work and leisure. All ecosystems of the world are potentially threatened by the effects of human activities. It is not unrealistic to forecast that, if human population will grow at the same rate as it does at present, all major ecosystems will be affected within another 100 years. If this would mean perturbation of major biogeochemical cycles and increased dominance of pests (fungi and insects), the necessity of worldwide ecosystem control would require enormous investments plus marked progress in the technology of cybemetics. It is not difficult to foresee that in that case only the people of countries with advanced technological knowledge will survive.

What is happening to the human population? Malthus expected food to be a major limiting factor, forecasting something like a J-shaped growth curve. Disasters like the Irish potato famine (after 1845) seem to support this viewpoint. However, in industrialized countries population growth tends to level off at present, showing something like an S-shaped growth curve. This growth retardation is mainly based on social considerations which, of course, can effect density-dependent reactions. Nevertheless populations in developing countries still grow exponentially.

Natural ecosystems are not yet adapted to our twentieth century way ofliving. It is questionable whether they will ever be. Ecosystems have at present many functions for the human population. These functions vary from production (most edible things are still derived from living organisms.) via "absorption" (ofpollution, of human-built structures) to regulation of the major biogeochemical cycles.

In conclusion, pollution, exploitation and environmental disruption cause instability of many ecosystems. The effects of human activities increase because human population still grows fast. In order to ensure a survival of the human population together with other organisms two things should be done. First, population growth has to be reduced as soon as possible. Second, the environment should be used in regard to conservation of functions rather than be exhausted.

Detailed information on environmental problems and the role of human activities can be found in the major textbooks on ecology (particularly Ehrlich et al. [11]) and in Detwyler [9], Murdoch [35], Moran et al. [32], Watt [54], Simmons [49], and Singer [50].

Acknowledgement

The author acknowledges the comments ofProf. Dr. L. Vlijm and Mr. L. C. Braat, and the typing of the manuscript by Ms. A. Jessurun and Ms. G. M. M. Simonis.

References

I. Allee, W.C.: Cooperation among Animals with Human Implications, Schuman, New York 1951 2. Boughey, A.S.: Ecology ofPopulations, 2nd ed., MacMillan, New York 1973 3. Clements, F .E.: Plant succession: analysis of the development of vegetation, Pub!. Camegie Inst.,

Wash. 242, 1-512 (1916) 4. Colinvaux, P.A.: Introduction to Ecology, Wiley, New York 1973 5. Collier, B.D., Cox, G.W., Johnson, A.W., Miller, P.C.: Dynamic Ecology, Prentice-Hall, Engle

wood Cliffs 1973


6. Connell, J.H.: Ecology 42, 133 (1961) 7. Darwin, c.: The Origin ofSpecies by Means ofNatural Selection, 1859 (reprinted by The Modern

Library, New York) 8. Daubenmire, R.F.: Ecol. Monogr. 22, 301 (1952) 9. Detwyler, T.R. (ed.): Man's Impact on Environment, McGraw-Hill, New York 1971

10. Dobben, W.H. van, Lowe-McConnell, R.H. (ed.): Unifying Concepts in Ecology, Junk, Den Haag 1975

11. Ehrlich, P.R., Ehrlich, A.H., Holdren, J.P.: Ecoscience: Population, Resources, Environment, Freeman, San Francisco 1977

12. Elton, C.S.: Animal Ecology, Sidgwick and Jackson, London 1927 13. Eyster, c.: Micronutrient requirements for green plants, especially algae, in: Jackson, D.F. (ed.)

Algae and Man, Plenum Press, New York 1964, pp. 86--119 14. Gause, G.F.: The struggle for existence, Hafner, New York 1934 15. Grinnell, J.: Amer. Nat. 51,115 (1917) 16. Hardin, G.: Science 131, 1292 (1960) 17. Haufe, W.O.: Int. J. Biometeor. 10, 241 (1966) 18. Holling, C.S.: Can. Entomol. 91, 293 (1959) 19. Hutchinson, G.E.: A Treatise on Limnology, Vol.l, Wiley, New York 1957 20. Kessel, B.: Condor 55, 49 (1953) 21. Kozlovsky, D.G.: Ecology 49,48 (1968) 22. Krebs, C.J.: Ecology; Harper and Row, New York 1972 23. Lack, D.: J. Anim. Ecol. 14, 12 (1945) 24. Leslie, P.H.: Biometrika 33, 183 (1945) 25. Leslie, P.H., Ranson, R.M.: J. Anim. Ecol. 9, 27 (1940) 26. Liebig, J.: Chemistry in Its Application to Agriculture and Physiology, Taylor and Walton, London

1840 27. Lotka, A.J.: Elements of Physical Biology, Williams and Wilkins, Baltimore 1925 28. MacArthur, R., Wilson, E.O.: The Theory ofIsland Biogeography, Princeton Univers. Press, Prin-

ceton 1967 29. McNaughton, S.J., Wolf, L.L.: General Ecology, Holt, Rinehart and Winston, New York 1973 30. McNeilly, T.: Heredity 23, 99 (1968) 31. May, R.M. (ed.): Theoretical Ecology, Blackwell, Oxford 1976 32. Moran, J.M., Morgan, M.D., Wiersma, J.H.: An Introduction to Environmental Sciences, Little,

Brown, Boston 1973 33. Morris, R.F. (ed.): The Dynamics ofEpidemic Spruce Budworm Populations, Mem. Entomol. Soc.

Can. 31 (1963) 34. Muller, C.H.: Bull. Torrey Bot. Club 93, 332 (1966) 35. Murdoch, W.W. (ed.): Environment: Resources, Pollution and Society, Sinauer, Stamford 1971 36. Odum, E.P.: Science 164, 262 (1969) 37. Odum, E.P.: Fundamentals of Ecology, 3rd ed., Saunders, Philadelphia 1971 38. Odum, E.P.: Ecology, 2nd ed., Holt, Rinehart and Winston, New York 1976 39. Odum, H.T.: Ecol. Monogr. 27, 55 (1957) 40. Odum, H.T.: Environment, Power and Society, Wiley, New York 1971 41. Pielou, E.C.: An Introduction to Mathematical Ecology, Wiley, New York 1969 42. Pontin, A.J.: J. Anim. Ecol. 30, 47 (1961) 43. Poole, R.W.: An Introduction to Quantitative Ecology, McGraw-Hill, New York 1974 44. Raunkiaer, C.: The Life Form ofPlants and Statistical Plant Geography, Clarendon Press, Oxford

1934 45. Ricklefs, R.E.: Ecology, Chiron, Newton 1973 46. Ricklefs, R.E.: The Economy of Nature, Blackwell, Oxford 1978 47. Sakagami, S.F., Fukuda, H.: Res. Pop. Ecol. 10, 127 (1968) 48. Shelford, V.E.: J. Morphol. 22, 551 (1911) 49. Simmons, I.G.: The Ecology of Natural Resources, Arnold, London 1974 50. Singer, S.F. (ed.): The Changing Global Environment, Reidel, Dordrecht 1975 51. Tansley, A.G.: The British Islands and Their Vegetation, Cambridge Univers. Press, Cambridge

1939 52. Volterra, V.: Nature 118, 558 (1926)


53. Watt, A.S.: J. Ecol. 35, 1 (1947) 54. Watt, K.E.F.: Principles of Environmental Science, McGraw-Hill, New York 1973 55. Went, F.W.: The Experimental Control of Plant Growth, Chronica Botanica, Waltham 1957 56. Whittaker, R.H.: Ecol. Monogr. 23, 41 (1953) 57. Whittaker, R.H.: Ecol. Monogr. 26, 1 (1956) 58. Whittaker, R.H.: Communities and Ecosystems, 2nd ed., MacMillan, New York 1975 59. Wingerden, W.K.R.E. van: Population Dynamics ofErigone arctica, Free University, Amsterdam

1977

Natural Radionuclides in the Environment

R. Fukai

International Laboratory of Marine Radioactivity, IAEA Principality of Monaco

Y. Yokoyama Centre des Faibles Radioactivites, CNRS-CEA Gif-sur-Yvette, France

Introduction

From the point of view of environmental research the significance of natural radioactive elements in the environment is twofold: one is the effects of ionizing radiation from these radionuclides to organisms living on the earth, especially man, and the other is the possible applications of natural radionuclides existing in situ as tracers for understanding geochemical processes which govern the distribution and fate of various pollutants in the environment. Although, of these two topics, the radiation effects may be more directly related to the health of the environment for man, this subject will be discussed only briefly in the present chapter, since it is practically impossible to reduce the natural radiation hazards by any means as long as man lives on the space-ship "Earth," and the present situation should be accepted. Thus, the scope of the chapter is focussed on the basis of the characteristics of natural radionuclides and their abundance in the environment to give a conci se review of how one can take advantage of their presence in situ for a better understanding of complicated processes taking place within the environment. Accordingly, the discussions which follow emphasize the introduction of a fundamental conept of time-scale measurements into environmental studies, rather than attempt a comprehensive literature coverage.

Characteristics of Natural Radionuclides

Since radionuclides are chemically similar to stable elements, that is, their outershell electron structures are not essentially different from those of stable elements,

48 R. Fukai and Y. Yokoyama

they undergo geochemical processes taking place in situ, such as dissolution, precipitation, sorption, complexation, etc., in the manner similar to those of stable elements. This is especially true for natural radionuc1ides coexisting with their stable counterparts in the environment. For example, radioactive 4°K behaves in situ exactly like stable 39K (isotopic abundance, "" 93 %) and 41 K (isotopic abundance, ",,7%), although its isotopic abundance amounts to only 0.01 %. Atoms of radionuc1ides are, however, characterized by the unstability of their nuc1ei. This unstability causes a radioactive atom to decay to another atom, either radioactive or stable, with a probability specific to the initial atom. In other words, the probability of the decay of a given radionuclide within a given time depends on the

,radionuc1ide involved and, thus, the half-time of the decay, which follows the exponentiallaw, becomes a constant for that specific radionuc1ide, regardless of the amounts of the radionuc1ide present. This half-time of the decay, traditionally known as a "half-life," is considered a primary physical property characterizing each radionuc1ide. The time-dependency of the radioactive decay gives the possibility that radionuc1ides could be used as unique tools for determining rate constants of various environmental events, that is, the time-scale of geochemical processes. The half-lives of natural radionuc1ides range from a fraction of a second to 1016 years, thus providing theoretically all choices for the application to different time-scale processes. In practice, however, only radionuc1ides of relatively longer half-lives (generally longer than, at least, several days) have been applied for this purpose.

In addition to the time-specificity, each radionuc1ide decays in a scheme specific to that radionuc1ide, emitting spontaneously particles and/or radiation with characteristic energies. This is another important characteristic of radionuc1ides which facilitates the identification of radionuc1ides contained in environmental materials with certainty. Due to the decay-scheme specificity and energy specificity, aided sometimes by the time specificity described above, most radionuc1ides can be identified and quantitatively determined at very low levels often found in complex environmental matrices. The accuracy of the radiological measurements suffers, in general, much less from reagent blanks and procedural contamination inherent in all trace measurements, which are especially critical for stable element measurements. In addition, high sensitivity and accuracy in radiological measurements can be attained by relatively modest laboratory facilities and instrumentation. For example, the combination of radiochemical separation procedures and successive cxspectrometry achieves the measurement of concentrations of transuranic elements - though they are not natural radionuc1ides - as low as 10 - 20 "" 10 - 22 gram of the elements per I gram of environmental matrices with associated errors within ±20% or less [1,2]. Thus, the higher sensitivity and accuracy of radiological measurements strengthen the utility of radionuc1ides as environmental tracers.

The characteristics of radionuc1ides described above apply not only to natural radionuc1ides but also to artificial radionuc1ides. The unique feature of natural radionuc1ides can, however, be identified among the interrelationships between the primordial actinide parents, 232Th, 235U and 238U and their respective descendant products. All of these parent nuc1ides decay in nature by emitting cx-partic1es (exc1uding spontaneous fission, the frequency ofwhich is much lower compared with that of cx-decay) to their respective daughter nuc1ides, all of them are again radio-

Natural Radionuclides in the Environment 49

active. Thus, relationships similar to those between the parents and the daughters are again established between the daughter and the grand-daughters. In analogy, the series of these relationships continues to be established until the last member of the series of the radioactive descendants decays to a stable nuclide. Since, practically, the loss of the mass of radioactive nuclei occurs only by the emission of aparticles during radioactive decay, the differences in mass numbers between the parents and their daughters are always either zero (in the case of ß-decay) or four (in the case of a-decay), the latter of which corresponds to the mass number of aparticles. This implies that the mass number difference between an actinide parent and its descendant nuclide is always a multiple of four. As the mass numbers of the parents 232Th, 235U and 238U can be expressed respectively as 4n, (4n + 3) and (4n + 2), where n represents arbitrary integral numbers, there is no possibility that a descendant nuclide characterized by its mass number is produced from a different initial actinide parent, although different isotopes of the same element are produced from different initial parents. For example, 226Ra is always a descendant from 238U and 228Ra from 232Th and never vice versa, although 226Ra and 228Ra are chemically identical. Thus, these three primordial actinide parents, 232Th, 235U, and 238U form respectively an independent series of their descendant nuclides characterized by their mass numbers. These series are called "radioactive decay chain" or "radioactive decay series."

In a closed system of sufficient age in the environment a descendant radionuclide in a given decay series exists in a steady state equilibrium, since the radioactive decay law dictates that the rate of production of the descendant from its immediate parent is equal to the rate of decay of the descendant itself to its daughter. In the dynamic environment such as the hydrosphere, however, the closed system is often disrupted by the separation of a daughter nuclide from its immediate parent, as the daughter may be chemically quite different from its parent. For example, the daughter 228Ra as a member of alkaline earth elements is much more soluble in ambient water than its immediate parents 232Th, so that 228Ra tends to dissolve into the liquid phase, leaving 232Th behind in the solid phase. As so on as a daughter radionuclide leaves the original closed system of a given decay series, it starts to decay in the open ambient system with its own half-life, producing its daughter and successive descendants, as its production is not any more supported by the decay of its immediate parent. In the original system, on the other hand, the regrowth of the missing daughter takes place after the reclosure of the system, the rate of regrowth being governed by the half-life of the missing daughter. Since the restoration ofthe steady state equilibrium takes approximately six half-lives ofthe daughter, the degree of disequilibrium observed meanwhile in the original system between the immediate parent and its daughter indicates a measure of the length of time between the separation of the daughter from the system and the observation. Thus, the quantitative information on the degree of disequilibrium in an actinide decay series gives the time-scale of the geochemical event, in this ca se the time elapsed since the reclosure of the system. Although the model process described above represents only one minor example among many approaches which have been made to the measurement of time-scales of various geochemical processes, it demonstrates clearly the effectiveness of the approach, which is not possible to achieve without the time specificity of radioactive decay and the characteristic fea-


ture of the natural actinide decay se ries described. The wider application of these approaches to environmental sciences is believed to assist the further understanding of the various complicated environmental processes.

Classification of Natural Radionuclides

More than 60 radionuclides are known to occur naturally in the environment. These are classified in two groups according to their origin: terrigenous and cosmogenie.

The terrigenous radionuclides are believed to have been present already in the rocks and minerals of the earth's ernst when the earth was formed, and include long-lived primordial nuclides coexisting with their stable element counterparts as weIl as the three primordial actinide parent nuclides, 232Th, 235U, and 238U, and their descendant products. Although stable isotopes of some members of the decay series also exist in nature, geochemical pathways of the radionuclides in the decay series in the environment are not necessarily similar to those of the corresponding stable isotopes, as their genetical histories are entirely different from those of stable isotopes.

The cosmogenic radionuclides, on the other hand, are produced continually in the earth's atmosphere by the bombardment of atoms of nitrogen, oxygen, argon, etc., with cosmic ray particles originating from outer space. They either are brought down to the earth's surface by precipitation and dry fallout or enter into geochemical processes taking place on the earth's surface in a gaseous phase. AIthough stable isotopes of all known cosmogenic radionuclides exist in nature, the degree of the isotopic exchange between cosmogenic radionuclides and their corresponding stable isotope depends very much on the forms in which these radio nuclides are introduced into the geochemical processes.

Examples of natural radionuclides belonging to the various categories stated above are given below, with brief descriptions oftheir characteristic features. Considering the applications of these radionuclides as geochemical tracers, only principal members having relatively long half-lives (normally longer than 100 days) are listed in each category, except for some radionuclides which are known to be useful for specific purposes.

Terrigenous Radionuclides

Primordial Nuclides Coexisting with Stahle Counterpart

In this category there are at least 14 radionuclides known at present to occur in nature. The common characteristic of these radionuclides is their very long halflives, which range from 107 years to 1015 years. Although their chemical properties vary widely, all of these radionuclides are considered to be weIl exchanged isotopically with their stable isotopes in the environment, even in the rocks and minerals of the earth's ernst, due to their primordial occurrence at the formation of the earth. Thus the environmental behaviour of these radionuclides is similar to that of their stable counterparts.

Natural Radionuc1ides in the Environment 51

The principal members in this group are 4°K and 87Rb. Both radionuclides belong to the alkali metal group, so that they are dispersed widely over the whole environmental spheres, especially in the hydrosphere, due to their susceptibility to weathering. The radioactivity of 40K thus represents over 90% of the total radioactivity of sea water, while that of 87Rb corresponds with approximately 1 %. In a closed system the decay of 4°K and 87Rb respectively to 40 Ar and 87Sr provides the basis of the most popularly-used methods for geochronology. The 4°Kro Ar method especially constitutes one of the most important geochronological methods, which can be applied to a variety of geochemical materials such as rocks, minerals, meteorites, manganese nodules, sediments and so forth. Although less widely applied, 87Rb can be used like 4°K. These geochronological methods are applicable for determining the ages ofmaterials ofthe order of, at shortest, 105 years or longer. The strict conditions of the closed system for their application, however, limits the expansion of their application to studies of dynamical geochemical processes.

Actinide Decay Series

This group includes three primordial actinide parent nuclides, 232Th, 235U, and 238U and their respective descendant nuclides. These descendant products consist ofabout 35 radioisotopes ofPb, Bi, Po, Rn, Ra, Ac, Th, Pa, U, etc. As described previously the three actinide parents form respectively an independent series of their descendants characterized by their mass numbers. Principal members ofthese decay series are schematically presented below. Considering the possible application of these nuclides as geochemical tracers, only those having half-lives longer than 100 days, except for 222Rn (half-life: 3.8 days) and 234Th (half-life: 24.1 days), are given in these schemes. In the following schemes solid arrows designate the direct decay from a parent to its daughter, while broken arrows represent the decay from a parent through short-lived intermediate nuclides to its· descendant. The half-lives and types of decay indicated in the schemes refer to the parent ofthe specific decay listed. Thorium decay series (mass numbers: 4n):

232Th 1.4 x10 10 year ) 228Ra ~.2..r:~-+228Th --.!:?.1'~':...-+208Pb(stable). a ß a

Actinium decay series (mass numbers: 4n + 3):

235U -.7:!.:..t~8~~-+231Pa 3.4x10 4year ) 227Ac _2..!.~~a.':.--+207Pb(stable). a a a. ß

Uranium decay series (mass number: 4n+2):

238U 4.5 x10 9 year )234Th~4.1~~-+234U 2.5x10 5 year )230Th7.5x104year ) a p a a

226Ra 1622 year ) 222Rn ~.~~al..-+210Pb ~~r=~-+210pO 138 day ) 206Pb (stable). a a p a

As has been pointed out previously, the formation ofthe radioactive decay disequilibria in a given enclosed system provides the basis for various geochronological methods, especially for the estimation of sedimentation rates in the marine environment. One of the most widely used methods for this purpose is the measure-


me nt of excess 230Th or 231Pa in sediment cores [3-5]. As can be seen in the above schemes, 230Th and 231Pa are the daughters Of 234U and 235U respectively. While uranium is soluble in sea water forming a carbonate complex, 230Th and 231Pa are hardly soluble in sea water and tend to settle to the sea bottom after having been scavenged by sinking particles. The equilibria within the decay series in the lithogenous fraction of the sediment are thus distributed by the addition of the unsupported excess of 230Th or 231 Pa. If the rate of sedimentation has been constant over the period given and also if there has been no significant post-depositional migration of deposited isotopes, the excess of 230Th or 231 Pa measured at a depth of the sediment core can be used as a measure ofthe time interval between the deposition and observation. This method can be applied to a time-scale range of the order of 105 years. Although there are severallimitations, a similar principle can be applied to the dating of authigenic marine deposits by measuring excess 234U with respect to 238U [6]. On the contrary, if uranium in sea water is specifically incorporated into calicareous materials such as coral, etc., leaving 230Th and 231Pa behind, then the calicareous material can be dated by the measurement of the degree of depletion of 230Th or 231Pa from the decay equilibria in the material.

Due to the long half-lives of the radionuclides involved, the above-mentioned examples of dating are only applicable to long time-scaled geochemical processes. In relation to the distribution and fate of various pollutants introduced in the environment, the much shorter processes are often more important. Applying the principle similar to that for long-lived radionuclides, the methods for measuring rate constants of shorter geochemical processes have been developed by utilizing medium half-lived nuclides, such as 210Pb [7-9], 228Th [10], 228 Ra [11], etc. The half-lives of these are in the range 1 ~ 21 years and are suitable for studying geochemical processes in a time-scale range from several years to 100 years. In addition to these applications, much shorter-lived radionuclides, 222Rn and 234Th, have also been used respectively for studying intensity of bottom current in the deep sea [12] and for estimating retention time of particulate matter in the surface mixed layer [13]. These examples indicate that possible application of shorter-lived radionuclides for understanding various geochemical processes is still an open research area where original ideas for specific applications are especially needed.

Cosmogenic Radionuclides

At least 14 radionuclides are known to be produced in the earth's atmosphere by nuclear reactions between gaseous atoms of the earth's envelope and cosmic-ray particles, such as high-energy protons, etc. The fluctuations in the rate of production of these radionuclides have been estimated from the measurement on meteorites, etc. These fluctuations are found to be within a factor of about two, if the averaged values over periods of their half-lives are taken, although for some radionuclides a significant increase in production rates some 106 years ago has been postulated. The fairly constant rates of production of these radionuclides make it possible to use them as unique geochemical tracers for studying various environmental processes, such as water mixing, particle transport, sedimentation, etc. Since a steady state source term of the inputs of these radionuclides into the


earth's environment can be assumed, differences in their concentrations among various parts ofthe environment can be re1ated to the rates of geochemica1 processes taking p1ace in situ, taking into account the decay after input. A wide half-life range of these radionuclides also makes an appropriate choice possible of radionuclide to be used depending on the time-scale of geochemica1 processes involved.

The prinicipal members of geochemical importance in this group inc1ude 3H (half-life: 12.3 years), 7Be (half-life: 53 days), lOBe (half-life: 2.5 x 106 years), l4C (half-life: 5,730 years), 26 Al (half-life: 7.4 x 105 years) and 32Si (half-life: ",700 years). Among these radionuclides, considerable amounts of 3H and l4C have been artificially produced by the recent nuclear explosion tests and added to the earth's environment. In addition to the naturally-occurring global inventories of 3H and l4C, representing respectively 34 MCi and 300 MCi [14], the excess inventories of these radionuclides due to a number of the explosion tests conducted mainly during the 1960s are estimated to amount to approximately 4,500 MCi for 3H [15,16] and 6 MCi for l4C [17] up to the early 1970s. Thus, it has become necessary to take into account the presence ofthe artificially injected 3H and l4C in the environment, when they are to be used as tracers of various geochemical processes.

Depending on their half-lives, these radionuclides have been widely used for studying various processes which take place in the environment. For example, the short-lived 7Be has been used as a tracer for studying atmospheric fallout rates and surface sea water mixing [18] and the medium-1ived 3H for the exchange between the intermediate and deep-water masses in the marine environment [19,20], while longer time-scale processes such as deep-water circu1ation and sedimentation processes have been investigated with the aid oflong-lived l4C [21]. Although difficulties in measurement and the specific behaviour of the stable counterpart limit its application as a geochemical tracer, 32Si can be used to complement l4C in a similar way to cover shorter time-scale processes [22]. The very long ha1f-1ives of lOBe and 26 Al make them excellent tracers for geochrono10gical studies on deep-sea sediments [23, 24], although the analytica1 difficulties involved, especially for the positron emitter 26 Al, prevent their wider applications.

Abundance in the Environment

In order to obtain a general picture of their relative abundance in various parts of the environment, the concentrations of the principal natural radionuclides in the atmosphere, the marine environment and the earth's crust are given in Table 1. The compilation has been based on the values presented in various publications, especially those of Koczy and Rosholt (1962) [25], LaI and Peters (1967) [14], Joseph et al. (1971) [26], Cherry and Shannon (1974) [27], and Burton (1975) [28]. Since the purpose of the compilation is to give comparative ideas of the abundance of these natural radionuc1ides with respect to their geochemical applications, and not to present data sources for precise computations, the values included in the table

54 R. Fukai and Y. Y okoyama

only show their orders of magnitude due to the variabilities and uncertainties of available data.

In compiling Table 1, the atmospheric concentrations of the nuclides listed have been computed based on the data presented by LaI and Peters (1967) [14], taking the earth's surface area as being 5.1 x 108 km2 and the thickness of the atmosphere at 0 °C under 1 atm as being 8 km. The concentrations in sea water and marine sediments have been based mainly on the values given by Joseph et al. (1971) [26] with modifications. The concentrations in the earth's crust have been computed from the average concentrations ofvarious elements in the continental crust reported by Tayler (1964) [29], assuming the radioactive decay equilibria.

As has been stated above, Table 1 presents, for the purpose of comparison, the average concentrations of naturaIly-occurring radionuclides in the different geochemical spheres of the globe, namely, the atmosphere, the hydrosphere (represented by the sea) and the lithosphere (represented by the earth's crust). Thus, the variability of the concentrations, within an individual geochemical sphere, has been disregarded in compiling the table. It should be emphasized, however, that the variation of the concentration of a radionuclide may be quite large within a given geochemical sphere. In general, these concentration variations are less pronounced within the relatively homogeneous atmosphere or hydrosphere than in the heterogeneous lithosphere, although the variations may be very significant even in the more homogeneous geochemical spheres. For example, while a terrigenous radionuclide, 238U, is distributed fairly homogeneously everywhere in sea water, the situation is obviously quite different over the earth's crust; acidic igneous rocks like granite is known to contain more uranium as weIl as thorium (> 5 ppm of U and ",20 ppm ofTh) than sedimentary rocks such as limestone, sandstone, etc. (1-2 ppm ofU and ",2 ppm ofTh) [30]. This results in local differences ofthe content ofthese radionuclides in soil covering various parts ofthe earth's crust. Depending on the origins of soil, its radionuclide content varies often by a factor of 10, sometimes by a factor of 100. In a few cases, such as those in the coastal areas of Brazil and India, where the terrain is covered with monazite-rich deposits, the thorium content of these deposits have been found to be extremely high; the content of thorium in the monazite sand in Kerala region (India) reaches up to 10% by weight [31].

While the production rates of the cosmogenic radionuclides in the atmosphere are considered to be fairly constant for many thousands of years in the past, inhomogeneity of these radionuclides within the same geochemical spheres exists with respect to their specific activities as weIl as with respect to their absolute concentrations. For example, it has been weIl established that the specific activity of 14C in the air is decreasing recently due to the diluting effect of releases into the atmosphere of CO2 from the increasing fossil fuel burning during the present century (the "Suess effect"). Fossil fuels are practically 14C-free due to the decay of this isotope (halflife: 5,730 years) during the isolation from the earth's surface for a geological time-scale. By 1954 the specific activity of 14C in the atmosphere was estimated to be reduced by approximately 5% due to this effect [32]. In analogy, local specific activity differences of cosmogenic radionuclides such as 3H and 14C may be caused by similar dilution processes; in urban and industrial areas where fossil fuel burning (including that of automobiles) is intensive, the specific activities

Tab

le 1

. C

once

ntra

tion

s o

fnat

ura

l ra

dion

ucli

des

in t

he e

nvir

onm

ent

Rad

ionu

clid

e H

alf-

life

A

tmos

pher

ea

Mar

ine

envi

ronm

ent

Con

tine

ntal

ear

th's

Z

cr

ust

S m

g 1

-1

pC

il-

1 Se

a w

ater

M

arin

e se

dim

ent

..... :::..

:;0

pCi

1-1

pC

i kg

-1

pCi

kg

-1

po

pp

m

ppm

p

pm

0

- o· :;

<= T

erri

geno

us r

adio

nuc/

ide

[ 4°

K

1.25

x 1

09 y

4

.8x

lO-2

34

0 ~3

~2 x

104

2.

4 1.

7 x

104

Cl> '"

87R

b 4.

7 x

10'0

y

3.4

xI0

-2

3 25

2.

2 x

103

S·

Tho

rium

dec

ay s

erie

s S- Cl>

228R

a 6.

7 y

1 x

lO-

14 ~0.002

2x

10

-9

450

5 x

10

-9

1.1

X 1

03

tri

:;

228T

h 1.

9 Y

2

x 1

0-1

5 2

X 1

0-3

7

xlO

-1O

56

0 1.

4 x

10

-9

1.1

x 10

3 $

. .....

0

232T

h 1.

42 x

lO

'Oy

x 1

0-6

1

X 1

0-4

5

540

10

1.1

x 10

3 :;

3 A

ctin

ium

dec

ay s

erie

s Cl

> a 22

7Ac

21.6

y

<1

.5x

lO-

12

<0.

1 6

x 1

0-9

43

0 5.

5 x

10

-10

40

231P

a 3.

43 x

104

y

2 x

10

-12

I X

1

0-4

I

X 1

0-5

45

0 I

x 1

0-6

40

23

5U

7.13

X 10

8 y

2.1

X 1

0-5

0.

05

7 x

10

-3

15

0.02

40

Ura

nium

dec

ay s

erie

s 21

0Pb

21 y

6

x 1

0-1

3 0.

05

5x

10

-8

4x

103

1

X 1

0-8

90

0 21

OpO

13

8 d

7 x

10

-15

0.03

9

x1

0-

1O

4 x

103

2 X

1

0-1

0 90

0 22

2Rn

3.8

d ?

1.3

X 1

0-1

5 0.

2 3

x1

0-

11

5 x

103

<6

X

10

-12

<9

00

22

6Ra

1662

y I

x 1

0-1

0 0.

1 4

x 1

0-6

4

x 1

03

I X

1

0-6

90

0 23

0Th

7.52

x 1

04 y

2

X 1

0-1

1

4 X

10

-4

2 X

1

0-4

4

X 10

3 4

X 1

0-5

90

0 23

4Th

24.1

d

4 x

10

-14

1 1.

4 x

10-1

1

330

4 x

10

-11

900

234U

2.

48 X

lO

S y

2 X

10

-7

1.2

8 x

10

-5

500

1.5

x 1

0-4

90

0 23

8U

4.5

X 10

9 Y

3

X 1

0-3

1

300

3 90

0

Cos

mog

enic

rad

ionu

c/id

e 3H

12

.3 y

6

x1

0-

17

6x

10

-4

1.7

X 1

0-1

5 1.

6 x

10-2

7B

e 53

d

5x

lO-

19

2 X

1

0-4

<

5

x 1

0-

14

<2

0

lOB

e 2.

5 x

106

y 4

x1

0-

15

6 x

10

-8

3 X

10

-14

4.5

xlO

-7

(1-3

) x

10

-7

1-4

14C

57

30y

4x

10

-13

2

xlO

-3

2 X

1

0-1

1 0.

1 (0

.1-1

) x

10

-7 4

0-40

0 26

AI

7.4

x 10

5 y

4x

lO-

19

7x

10

-12

3 X

10

-16

5 X

1

0-9

32

Si

~700y

6x

lO-

19

IxlO

-8

6 X

10

-16

I x

10

-5

<2

x 1

0-1

0 <

3

V.

V.

a B

ased

on

air

at N

.T.P

., 0

°C a

nd

1 a

tm


of 14C is expected to be lower; around volcanos where the injections of juvenile H20 and CO2 occur, the isotopic dilution of both 3H and 14C takes place; and so forth.

In addition to the inhomogeneity in their specific activities, the absolute concentrations of some cosmogenic radionuclides vary also within the homogeneous geochemical spheres. For example, due to their relatively short half-lives, the concentrations of 3H and 7Be in sea water decrease rapidly from sea surface with increasing depth; the limited volume ofH20 on the land surface, available for diluting tritiated water, not only maintains the specific activity of 3H of fresh water on the land surface higher than that in sea water, but also elevates the absolute concentration of 3H in the fresh water relative to sea water [33]. Since the causes of these concentration gradients are usually related closely to some specific geochemical processes, the radionuclides involved can become useful tracers for studying these processes.

Radiation Effects

Although, as previously stated, the effects of ionizing radiation on natural ecosystems and man from naturally-occurring radionuclides are unavoidable hazards, the recent introduction of artificially-produced radionuclides into the environment as a result of human activities has focussed attention also on natural radiation effects. In order to assess the additional radiation effects from the man-made radionuclides in the environment, it is essential to know the precise effects of natural radiation as a basis. Despite the quantitative importance ofthe radiation dose received by man in the environment from naturaIly-occurring radio-nuclides, compared with that from artificial radionuclides, the effects of chronical exposure to low-Ievel natural radiation are so far not weIl understood. Detailed review and evaluation of the information obtained so far on the radiation effects of naturallyoccurring radionuclides on man have been carried out by the United Nations' Scientific Committee on the Effects of Atomic Radiation (1977) [34].

Radiation effects are usually considered from two aspects: the somatic effects and the genetic effects. One consensus so far obtained is that there are no evident somatic effects on man from natural radiation. The genetic radiation effects are much more difficult to assess.

The radiation dose received by man varies widely depending on natural conditions ofhis habitat as weIl as on the conditions ofhis life. Since, as has been stated previously, the content of terrigenous radionuclides in soil varies geographicaIly, the external radiation dose received by man varies also depending on where he lives. While the global average ofthe absorbed dose rate in air from terrestrial radiation is estimated to be 4.5 J.l rad h- 1 (40 mrad y-1) [34], the value of 130 J.l rad h -1 (1,140 mrad y-1) has been estimated for the monazite area ofKerala coast in India [31]. Due to the increase of cosmic radiation the radiation dose rate increases also with the height of the habitat. At about 3,000 m altitude the radiation dose rate due to cosmic radiation ('" 35 mrad y-1 at sea level) has been estimated to increase by a factor of 3 [35]. In addition to the differences in natural surroundings, the living conditions ofindividuals affect the radiation dose received by man. For

Natural Radionuclides in the Environment 57

example, the radiation exposure indoors depends on building materials of houses; due to the higher content of terrigenous radionuclide, bricks and granite give usually much higher radiation dose rate than that from cement, pIaster, wood, etc. While the differences in radiation dose rate due to the building materials are usually within a factor of 10, those by a factor of 100 are sometimes observed [34]. Other factors such as feeding habits of individuals, types of work he is engaged in, etc. affect the radiation exposure of man.

It has been made clear, on the other hand, that some organisms living in the marine environment are receiving a much higher natural radiation dose than organisms living on land, due to the accumulation of 210pO, the descendant nuclide of 238U, in some organs of their bodies [27]. Since the natural radioactivity of the marine environment has not changed significantly during the evolutionary history of marine organisms (about 109 years), studies of evolutionary history of marine organisms in connection with radiation effects at cellular levels in these organisms may throw some light on the subject of assessing the effects of chronie low-Ievel radiation from natural radionuclides.

In connection with the increase in radiation effects due to use of nuclear power, it may be worthwhile mentioning the release of natural radionuclides by fossil fuel power production. It has been estimated that, due to natural radioactive elements such as uranium, thorium, etc. contained in coal and emde oil, the increase in radiation dose received by man living in the environs of a power plant is higher in the case of a coal-fired power plant than a nuclear power plant, when compared on unit power production [36]. This aspect should be considered when the safety of power production is assessed.

Application of Geochemical Tracers

As the major scope of the present chapter is to give fundamental insight into the expanding of the use of natural radionuc1ides as geochemical tracers for studying various environmental processes, a brief summary of guidelines for applications is considered worthwhile to obtain further ideas on the possibilities of the applications and their limitations. Examples ofthese applications are given below, being classified by the environmental processes involved. These examples are not exhaustive, but an attempt is made to describe the fundamental ideas on which they are based, rather than go into technical details.

Transport Processes

The processes in this category include mainly those of transport of particuiate matter, such as fallout particles, dusts, aerosols, lithogenic and biogenie particles, etc., in the atmosphere as weIl as in the aquatic environment. Since the time-scale of these processes ranges from days to years, applicable tracers should have comparable half-lives. In addition to an appropriate half-life, a suitable tracer should be particle-forming or easily sorbable. In order to fulfil these requirements, 7Be (half-life: 53 days), 210Pb (half-life: 21 years),

58

22BRa (half-life: 6.7 years), 22BTh (half-life: 1.9 years), and 234Th (half-life: 24.1 days)

R. Fukai and Y. Yokoyama

are considered suitable depending on the time-scale of the processes to be studied. Though not listed in Table I 224Ra (half-life: 3.6 days), the daughter of 22BTh, may also be used for this purpose for very short time-scale measurements in some specific conditions. In considering the application of a tracer, the chemical property of the tracer itself as weIl as that of its immediate parent and daughter should be taken into account. For example, it is important to note when applying isotopes ofthorium and radium as tracers that thorium isotopes are insoluble in the aquatic environment, associating themselves easily with particulate matter, while radium isotopes are more soluble, tending to be incorporated in phytoplankton.

Mixing Processes

Since suitable tracers for mixing processes in the atmosphere or in the hydrosphere have to be mixed homogeneously within an air mass or a water mass, they should be in gaseous form for atmospheric mixing or in easily soluble form for water mass mixing. While the time-scale of atmospheric mixing processes, in general, is fairly short, ranging from days to months, water mixing processes proceed much more slowly taking from months to some 1,000 years.

Although cosmogenic radionuclides such as 3H as tritiated water vapour and 14C as carbon dioxide gas are in suitable forms as tracers for atmospheric mixing, their applications for this purpose are limited due to their too long half-lives compared with the time-scale of atmospheric processes and to the characteristic geographical pattern oftheir production. On the contrary, the bomb-produced 3H and 14C can be used widely as tracers for air-mass transport and mixing, because of their localized production pattern. To a limited extent the short-lived gas 222Rn (half-life: 3.8 days) may be used, together with its descendants 210Pb and 210pO, for tracing short-term mixing between land air-mass and marine air-mass, since the content of 222Rn released in the air above land is believed to be higher than that in the marine atmosphere.

As already described, 3H and 14C serve as excellent tracers for water mixing processes of various time-scales. While 3H is chiefly used for tracing the exchange between intermediate water and deep water, 14C can be applied to long-term deepwater circulation. U nder specific conditions 32Si (half-life: '" 700 years) may be used in a similar manner to complement these two isotopes. F or shorter time-scale processes such as surface water mixing, etc., short-lived 7Be, remaining soluble in sea water for a certain length of time, can be applied.

Sedimentation Processes

In a broad sense the sedimentation processes include the growth of manganese nodules in the deep-sea bottom, the growth of coral in tropical areas, etc. The tracers to be used for studying sedimentation processes are required to be incorporated in depositing particles under specific conditions. Most sampies ofthese studies


are usually deep-sea sediment cores representing a time-scale range of 1 04 ~ 106 years; the half-lives of radionuclides suitable for this purpose are normally long. U p to now, the widest applications of natural radionuclides have been made in this research field, geochronology, taking advantage of the characteristics of natural radionuc1ides previously described. For long-term sedimentation processes, wide use of long-lived nuclides from the actinide decay series such as 226Ra, 230Th, 232Th, 231Pa, 234U, 238U, etc., as weIl as cosmogenic nuclides such as lOBe, 26 Al, and 32Si, has been made.

Since sedimentation processes are directly related to the transport of depositing particles to the bottom, the tracers suitable for short-term partic1e transport processes, such as 7Be, 2l0Pb, 228Ra, 228Th, 234Th, etc., should be useful also for studying sedimentation processes in shallow water.

Exchange Processes

The processes c1assified in this category are characterized by the mass transfer of material through an interface, such as the air-water interface and water-sediment interface. Thus, in principle, volatilization or dissolution of a radionuc1ide through an interface makes a quantitative estimation ofthe exchange flux possible when the concentration difference of the nuc1ide between two phases is known. The release of 222Rn from deep-sea sediments is an example of such application, and the thickness of the 222Rn-bearing water layer has been used for estimating slow current velocity near the sea bottom. Soluble radionuc1ides such as 226Ra, 228Ra, etc., may be used for similar purposes.

Pathway Indicators

The difference in the 234U F38U activity ratio between sea water and lithogenous particles has been applied for estimating the pathway ofuranium taken up by marine organisms [37]. While the 234U F38U activity ratio of dissolved uranium in sea water has been weIl established as 1,14±0.02, the ratio in lithogenous suspended matter in sea water should be c10se to one due to the radioactive decay equilibrium. Thus, by determining the 234U /238U activity ratio in marine organisms, the pathway of the up-taken uranium, that is, whether it is taken-up from sea water or through particulate ingestion, can be estimated. Although not frequently applied so far, the similar type of approach to various environmental uptake problems is considered to have wide application potential.

References

1. Fukai, R., Ballestra, S., Holm, E.: Nature 264, 739 (1976) 2. Holm, E. et al.: Oceanol. Acta 3, 157 (1980) 3. Goldberg, E.D., Koide, M.: Science 128, 1003 (1958) 4. Goldberg, E.D., Griffin, J.J.: J. Geophys. Res. 69, 4293 (1964) 5. Griffin, J.J. et al.: Deep Sea Res. 19, 139 (1972) 6. Ku, T.-L., Broecker, W.S.: Deep Sea Res. 16,625 (1969)


7. Goldberg, E.D. et al.: Geochem. Jour. 10, 165 (1976) 8. Goldberg, E.D. et al.: Geochim. Cosmochim. Acta 42, 1413 (1978) 9. Benninger, L.K.: Geochim. Cosmochim. Acta 42, 1165 (1978)

10. Koide, M., Bruland, W., Goldberg, E.D.: Geochim, Cosmochim. Acta 37,1171 (1973) I!. Moore, W.S., Krishnaswami, S.: Earth planet. Sci. Lett. 15, 187 (1972) 12. Schink, D.R. et al.: Trans. Nucl. Sci., N.S. 17, 184 (1970) 13. Bhat, S.G. et al.: Earth Planet. Sci. Lett. 5, 483 (1969) 14. Lai, D., Peters, E.: Cosmic ray produced radioactivity on earth. In: Encyc10pedia of physics, pp.

600-601, Springer-Verlag, New York 1967 15. Schell, W.R. et al.: World distribution of environmental tritium. In: Physical Behaviour of Radio

active Contaminants in the Atmosphere, pp. 375-395, IAEA, Vienna 1974 16. Michel, M.L.: Nature 263, 103 (1976) 17. Fairhall, A.W. et al.: Radiocarbon in the sea. In: Health and safety Laboratory Fallout Program

Quarterly Report, HASL-242, pp. 1-35 to 1-78, New York 1971 18. Silker, W.B.: Earth Planet. Sci. Lett. 16, 131 (1972) 19. Rooth, c.G., Ostland, H.G.: Deep Sea Res. 19,481 (1972) 20. Roether, W., Münnich, K.O.: Earth Planet. Sci. Lett. 16, 127 (1972) 21. Ribbat, B., Roether, W., Münnich, K.O.: Earth Planet. Sci. Lett. 32, 331 (1976) 22. Lai, D., Somayajulu, B.L.K.: Possible application of cosmogenic silicon-32 for studying mixing in

near coastal waters. In: Isotope Marine Chemistry (Goldberg, E.D., Horibe, Y., Saruhashi, K. eds.), pp. 145-156, Uchida-Rokakuho Publ. Co., Tokyo, Japan 1980

23. Guichard, F., Reyss, J.-L., Yokoyama, Y.: Nature 272, 155 (1978) 24. Raisbeck, G.M. et al.: Earth Planet. Sci. Lett. 43, 237 (1979) 25. Koczy, F.F., Rosholt, J.N.: Radioactivity in oceanography. In: Nuc1ear Radiation in Geophysics

(Israel, H., Krebs, A. eds.), pp. 18-46, Springer-Verlag, Berlin 1962 26. Joseph, A. et al.: Sources of radioactivity and their characteristics. In: Radioactivity in the marine

environment (Seymour, A.H. ed.), pp. 6-41, Nat. Acad. Sciences, Washington, D.C. 1971 27. Cherry, R.D., Shannon, L.V.: Atom. Ener. Rev. 12, 3 (1974) 28. Burton, J.D.: Radioactive nuc1ides in the marine environment. In: Chemical oceanography, vol. 3

(Riley, J.P., Skirrow, G. eds.), pp. 91-191, Academic Press, London, New York, San Francisco 1975

29. Tayler, S.R.: Geochim. Cosmochim. Acta 28, 1273 (1964) 30. Adams, J.A.S.: Radioactivity of the lithosphere. In: Nuc1ear Radiation in Geophysics (Israel, H.,

Krebs, A. eds.) Springer-Verlag, Berlin 1962 31. Gopal-Ayengar, A.R. et al.: Current Status of investigations on biological effects of high back

ground radioactivity in the monazite bearing areas of'Kevala coast in South-West India. In: Proc. Intern. Symp. on Areas of High Natural Radioactivity, pp. 19-28, Pocos de Caldas, Brazil 1975

32. Bascastow, R., Keeling, C.D.: Atmospheric carbon dioxide and radiocarbon in the natural carbon cyc1e. In: Carbon and the Biosphere, US-AEC Rep. CONF-720510, pp. 86-135, Washington, D.C. 1973

33. Kaufman, S., Libby, W.S.: Phys. Rev. 93, 1337 (1954) 34. United Nations Scientific Committee on the Effects of Atomic Radiation: Sources and effects of

ionizing radiation, 1977 report to the General Assembly, United Nations, New York 1977 35. Folsom, T.R., Harley, J .H.: Comparison of some natural radiations received by selected organisms.

In: The Biological Effects of Atomic Radiation, pp. 28-33, NAS-NRC, Publ. No. 551, Washington, D.C. 1957

36. Okamoto, K.: Fifth Intern. Congr. Internat. Radiation Protection Ass. Jerusalem 1980 37. Hodge, V.F., Koide, M., Goldberg, E.D.: Nature 277, 206 (1979)

The Nitrogen Cycles

R. Söderlund

Department of Meteorology, University of Stockholm S-10691 Stockholm, Sweden

T. Rosswall

SCOPE/UNEP International Nitrogen Unit Royal Swedish Academy of Sciences, Box 50005 S-10405 Stockholm, Sweden

Introduction

The purpose of this paper is to review the biogeochemical nitrogen cycle. The importance of nitrogen derives mainly from its fundamental nutritional role, but is also due to the fact that nitrogenous substances are important environmental pollutants. The possible effect of the interaction between nitrogen oxides and the stratospheric ozone layer is aglobai concern. There are also environmental problems, of local to regional character, in relation to nitrogen that deserve attention, such as the contamination of ground water by heavy applications of nitrogen fertilizer, thus making drinking-water from these sources unfit for human consumption. Increased levels of nitrogen in air and precipitation are creating changes in the atmospheric chemical climate, one result of which being eutrophication of the environment. The acidification of rain-water by nitrogen oxides is another example of the environmental consequences of nitrogen use.

The approach taken in describing the biogeochemical nitrogen cycle is based on a mass-balance study, primarily for the global system. The construction ofmass balances is useful for several reasons, one being that they constitute a conceptual model of a real system and can therefore serve as a useful organizational too1.

Even a crude mass balance enables one to evaluate information needs and data gaps, therefore allowing a better monitoring or data-gathering programme. In a complex system, a mass-balance model can provide important insight into the interactions between compounds that at first may seem unrelated or only loosely related. From a practical viewpoint, mass-balance studies can be used to evaluate the importance of anthropogenie versus natural sources of flows of a substance in a given system.

62 R. Söderlund and T. RosswalI

Examples from subsystems will be given to underline the heterogeneity of the global systems.

Basic Chemical Considerations

Nitrogen can occur in both inorganic and organic forms. The inorganic forms occur in valence states from - III to + V. The compounds possess a wide variety of chemical and physical characteristics, ranging, for example, from highly acidic to mainly basic in the natural environment.

The different chemical forms of nitrogen discussed in this chapter are listed in Table 1. Their relation to water - the global solvent and carrier of many chemicals - is also given, since this is important in understanding the biogeochemical cycle of nitrogen.

Owing to their interconvertibility in the atmosphere, the two nitrogen oxides nitric oxide (NO) and nitrogen dioxide (N02) are given the joint designation NOx •

Chemical Transformations of Nitrogen Compounds in the Environment

Introduction

The biogeochemical nitrogen cycle is unusually complex in that it involves a large number of compounds, ranging from the reduced forms (e.g., NH3) to the highly oxidized forms (e.g., N03"). Most of the transformations involved in the biogeochemical nitrogen cycle are carried out by microorganisms in water and soil. Figure 1 shows a highly schematic nitrogen cycle focusing on the biological reactions. Molecular dinitrogen gas can be fixed by certain microorganisms (biological nitrogen fixation), either free-living or in association with certain plants (process 1; Fig. 1). Nitrogen in living organisms occurs mainly in various organic compounds, e.g. proteins, amino acids and nucleic acids. Through mineralization (2) (also called ammonification), organic nitrogen is transformed to inorganic ammoniumnitrogen. This process is mainly carried out by microorganisms (bacteria and fungi), and the NHt formed is readily available as a nitrogen source for most organisms by incorporation into new biomass (immobilization; 3). Ammonia can be further oxidized to nitrite (N02) and nitrate (NO;) by certain microorganisms (4) through the process of nitrification. Nitrate is also a suitable nitrogen source for many organisms, and one of the most important for plants.

A number of microorganisms can reduce nitrate, either to ammonium (nitrate assimilation, 5, or non-assimilatory nitrate reduction, 6) or via nitrite to nitrous oxide (N20) or nitrogen through denitrification (7). N20 can also be produced during nitrification.

Nitrogen Fixation

Nitrogen fixation can be carried out by a number of microorganisms, either freeliving or in association with certain plants. The fixation of nitrogen requires a high energy expenditure, and the enzymes (nitrogenase system) are highly susceptible to

The Nitrogen Cycles

Table 1. Chemical properties of some compounds discussed in the text

Compound

Gases Ammonia

Molecular nitrogen Nitrous oxide

Nitric oxide

Nitrogen dioxide

Acids Nitrous acid Nitric acid

Bases Amines

Others Peroxyacyl nitrate

Formula

N 2

N 2 0

NO

N02

HN02

HN03

R 3N

Properties

Colourless gas, purgent odour, highly soluble in water, basic reactions

Colourless gas, low reactivity Colourless gas, small reactivity at ambient

temperatures Colourless gas, low solubility in water,

reactive Reddish-brown gas, highly reactive,

reacts with water

Weak acid, source of the N02" ion Strong acid, forms soluble salts with

many metals

Depending on the nature of the substituent (s) R, amines have different characteristics

For R 1 = -OH:R2 =R3 =H gives hydroxylamine

water-soluble, irritating to the eyes, toxic compound

63

Valence state

-III

o +1

+II

+IV

+III +V

Fig.1. Schematic view of the biotic part of the nitrogen, biogeochemical cycle. 1. Nitrogen fixation, 2. Mineralization, 3. Immobilization (ammonia assimilation), 4. Nitrification, 5. Immobilization (nitrate and nitrite assimilation), 6. Non-assimilatory nitrate reduction, 7. Denitrification

oxygen. These faetors are probably the two most important whieh have restrieted the evolution of a nitrogen-fixing eapaeity to eomparatively few organisms of symbiotie assoeiations. Table 2 lists the most important nitrogen-fixing systems.

Free-living heterotrophie baeteria eapable of fixing nitrogen are eommon in most habitats, but their eapaeity to fix nitrogen is less than that of eyanobaeteria

64 R. Söderlund and T. Rosswall

Table 2. Examples of organisms known to fix dimoleeular nitrogen. Not all speeies of the genera mentioned neeessarily fix N2• Only examples are given; for a more eomplete list of nitrogen-fixing organisms, see [5, 34, 49, 50]

A. Free-living heterotrophie baeteria a) Aerobic

b) Anaerobie

B. Free-Iiving autotrophie baeteria a) Photo-autotrophie

b) Chemo-autotrophie e) Cyanobaeteria (blue-green algae)

C. Obligate symbiosis a) Root-nodule symbiosis

b) Symbiosis with eyanobaeteria Liehens

ii Liverworts

iii Pteridophytes iv Angiosperms

D. Assoeiative symbiosis

Azotobacter Beijerinckia Azospirillum Klebsiella Clostridum Desulfovibrio M ethanobacillus

Rhodopseudomonas Rhodospirillum Thiobaci llus Anabena Nostoe Gleotrichia Calothrix Trichodesmium

Rhizobium - Leguminosae Parasponia andersonii (formerly Tremai-cannabina) Actinomycetes - Ainus glutinosa (alder) Actinomycetes - Myria gale (bog myrtle) Actinomycetes - Hippophae rhamnoides (sea buekthorn) Actinomycetes - Casuarina equisetifolia

Nostoe - Peltigera Nostoc - Collema Nostoe - Blasia N ostoc - Anthoceros Anabena - Azolla N ostoc - Gunnera

a) Azobacter papsali with Papsalum notatum b) AzospiriIlum lipoferum with Digitaria decumbens

(blue-green algae) and symbiotic associations. The cyanobacteria can obtain energy through photosynthesis and often have a very efficient nitrogenase system; these organisms can be very important for the nitrogen economy of certain terrestrial ecosystems, such as tundra [24, 42] and rice fields [53].

Cyanobacteria do not only occur free-living but also in symbiosis with fungi in lichens and with the water fern Azolla. Nitrogen-fixing lichens may be very important for such systems as tundra, while the Azolla symbiosis could be of great significance in, for example, rice fields.

The economically most important nitrogen-fixing system is that between bacteria of the genus Rhizobium and plants of the Leguminosae family, such as clover (Trifolium ssp.), soy bean (Glycine max) , broad bean (Viciafaba), cow-pea (Vig-

The Nitrogen Cyc1es 65

na sinensis) , chick-pea (Pisum sativum) and lucerne (Medicago sativa). The use of these crops has proved to be of great importance - especially in areas where access to nitrogen fertilizers is sm all - in obtaining good harvests and sustaining soil fertility.

In recent years, the associative symbiosis between the bacterium Azospirillum lipoferum and the roots of plants such as the tropical grass Digitaria decumbens, or maize has received considerable attention [6, 18]. In a similar manner, the tropical forage grass Paspalum notatum seems able to support an active population of nitrogen-fixing Azotobacter paspali in the root zone [16, 17].

Mineralization and Immobilization

Organic bound nitrogen is mineralized to ammonium nitrogen mainly through the action ofmicroorganisms, although in the case ofcertain animals (e.g., protozoa, nematodes and earthworms) ammonium mayaiso be an excretory product. Bacteria and fungi use a range of enzymes, such as proteases, peptidases and deaminases, to produce NHt -N from organic nitrogen compounds. During the mineralization of organic matter, the microorganisms obtain energy through the oxidation ofthe organic compounds to CO2 . Since carbon is lost through respiration, nitrogen will eventually be present in excess and be liberated by the microorganisms to the environment, where it can be utilized, by for example, plants for growth, thereby again being immobilized into an organic form. Ammonium nitrogen can also be immobilized by microorganisms, especially if the carbon (energy) source is in abundant supply.

Nitrification

Certain autotrophic bacteria - nitrifying bacteria - can use inorganic nitrogen oxidation as an energy-yielding process. The genera Nitrosomonas, Nitrosococcus, Nitrosospira, Nitrosovibrio and Nitrosolobus oxidize ammonium to nitrite according to the following overall reaction:

~F' -65 kcal,

while Nitrobacter and Nitrococcus spp. oxidize nitrite to nitrate according to:

~F' -18 kcal.

These reactions can only occur in the presence of oxygen. The oxidation of N0"2 is generally more rapid than that of NHt , and nitrite rarely accumulates in the environment.

Ammonium-N can, however, also be oxidized to a variety ofcompounds, such as hydroxylamine, nitrite and nitrate by certain heterotrophic bacteria, yeasts, actinomycetes and fungi [22]. Nitrous oxide can be released as a by-product in autotrophic nitrification. The steps involved in ammonium oxidation by Nitrosomonas and also the possibility of nitrite reduction by the same bacterium are given in Fig.2.


-3 -1 0 +1 +2 +3 15 NHt -- 15NH20H _ [15~OH .... ;> 15.15N202H2] ----;.15NO _15N02-

: :

.. ...

15.15N20

14.15N202H2·· .. ~ 14.15N20 ~

14.14N20 ~

14NHt <- - - 14NH20H <---[14NOH ... ;> 14.14N202H2] _14NO ___ 14N02

Fig.2. Inorganic nitrogen metabolism in Nitrosomonas. Enzymatic reactions -+, possible enzymatic reactions _ ... , chemical reactions >. (From Ritchie and Nicholas, 1972 [36]).

The nitroxyl radical [HNO] is not stable in neutral aqueous solutions, being immediate1y dimerized to hyponitrous acid (H2N20 2). The possible production of N 20 under aerobic conditions may explain results obtained in denitrification studies, in which N 20 - earlier thought to occur only as one end-product of denitrification (see below) - has been detected as being produced under aerobic conditions [4]. The possible advantages for Nitrosomonas of possessing a nitrite-reducing pathway may be that they can detoxify nitrite if, for some reason, this fairly toxic compound accumulates, but that they mayaiso be able to use nitrite as a terminal e1ectron acceptor for survival in brief periods of anaerobic conditions [40]. It is important to keep in mind the possibility of N 20 production during aerobic conditions when discussing the global rates ofN 20 production and its possible impact on the stratospheric ozone layer.

Denitrification and Nitrate Assimilation

Under an aerobic conditions, certain aerobic bacteria can utilize nitrate in place of oxygen as a terminal electron acceptor, reducing it to the gaseous end-products NO, N20 or N2.

The main steps in denitrification are as folIows:

NO; ~N02 ~NO~N20~N2'

Dinitrogen gas and nitrous oxide are the two most important end-products, N 2 usually being produced in excess ofN20. Nitric oxide (NO) can, however, also be produced under certain conditions [48]. The use of nitrate as a terminal electron acceptor (anaerobic respiration) generates nearly as much energy as the aerobic respiration, and much more than the common fermentative pathways.

It is possible that there is a dose coupling between the metabolie pathways of nitrification and denitrification (Fig. 3), not only because ofthe fact that N20 constitutes one end-product for both reactions, but also in the light of the suggestion that the nitroxyl radical [HNO] is a necessary intermediate for the dimerization, when nitric oxide (NO) is reduced to nitrous oxide (N20), as suggested by De1-wiche and Bryan [15].

The Nitrogen Cycles 67

Fig. 3. Possible couplings between nitrification and denitrification. ~ denitrification, - - .. nitrification, _ ..... nitrate/nitrite assimilation, non-assimilatory nitrate reduction

Nitrate can be reduced not only to N 20 and N 2 during nitrate respiration, but it can also be reduced by an assimilatory pathway first being converted to NHt and then being incorporated into cell biomass in organic compounds. Certain microorganisms can also reduce N03 to NHt by a dissimilatory pathway under anaerobic conditions, and this, together with denitrification, is an energy-generating process [52].

Abiotic Nitrogen Transformation

A large number of compounds react with the nitrogen species found in the environment. Photochemical reactions involving, for example, ozone, and the fixation of dinitrogen - during combustion to form nitrogen oxides, and in the industrial manufacture of ammonia, mainly for fertilization purposes - are important abiotic reactions which determine the amount and forms of nitrogen found in the environment.

The photochemistry of the troposphere develops round the reactive oxygencontaining species. Light triggers the reaction whereby ozone and nitrogen dioxide are photodissociated and atomic oxygen is formed. This is followed by a large number of reactions which involve organic compounds and which are additional sources of reactive oxygen species. Among the compounds formed in this complicated reaction scheme are nitric acid and toxic compounds, such as PAN (peroxyl acyl nitrate).

Nitric acid is formed through the reaction of hydroxyl radicals and nitrogen dioxide:

OH + N02 --+HN03 (g).

The hydroxyl radical is formed in the reaction of atomic oxygen and water vapour, according to:

O+H20--+20H.

At elevated temperatures, nitrogen and oxygen combine to form nitric oxide. The reaction can be schematically written as:

-iO+-iN2 --+NO LlG = 866.6kJmol- 1; logK=9.82 at 278°K.

The yield of nitric oxide from this gas-phase reaction is dependent on both the reaction temperature and the detailed cooling process of the reactants. At a slow decrease in temperature of the gases, the dinitrogen and dioxygen are re-formed and the yield of nitric oxide is thus small.


In lightning, the fixation process is believed to involve the conversion ofN 2 and O2 into NO by high-temperature chemical reactions [56]. The air around and in the discharge channel is heated to temperatures in excess of 4000 K and then cooled, causing stabilization of the nitric oxide formed. Owing to the convective character ofmost thunderstorms and the poor solubility ofthe nitric oxide formed, the nitrogen oxide remains in the atmosphere for aperiod of the order of days before it is converted into more reactive forms (mainly nitric acid), which are susceptible to the deposition processes.

Most nitrogen fixation on the industrial scale is carried out by the Haber process. A mixture of N 2 and H2 is combined to NH3 at temperatures ranging from 675-900 K and at pressures of 100-150 times the atmospheric pressure. A catalyst containing iron and aluminium is used to enhance reaction rates under these conditions. The mixture of dinitrogen and hydrogen is obtained from re-forming of natural gas or from coke-oven gas.

A large part of the ammonia produced in the world is used for fertilizer manufacture. Ammonium nitrate is produced by allowing ammonia to react with nitric acid,

The nitric acid is obtained industrially in the Ostwald process from the oxidation of ammonia by oxygen at elevated temperatures and by the use of a platin um/ rhodium catalyst.

Global Inventories of Nitrogen

Introduction

The· three compartments - atmospheric, aquatic and terrestrial systems - will be used when describing the global nitrogen cycle in accordance with Söderlund and Svensson [47]. In defining a system, boundaries must be assigned to the compartments. The boundaries are by no means obvious, and hence often have no physical meaning, the relevance only lying in relation to the description of the process under study. For example, the soil atmosphere is regarded as part of the atmosphere when nitrogen fixation is considered. On the other hand, when nitrous oxide is evolved via denitrification, it is not considered as part of the atmosphere until it has escaped from the surface of the soil.

The global inventories of nitrogen for the terrestrial, aquatic and atmospheric systems are given in Table 3. A large quantity of nitrogen is found in sedimental deposits and rocks (574.108 Tg [51]). This reserve of nitrogen is not available to the biosphere and is not further considered here.

The inability of most living species to utilize the molecular form of nitrogen gives rise to the paradox that while nitrogen is one of the most abundant elements in the biosphere it is considered as being in limited supply. The reason for this is that 99.96% ofthe nitrogen is in the form ofN2 , with only 0.04% in a combined form.

The Nitrogen Cyc1es

Table 3. Global inventories of nitrogen in the terrestrial, oceanic and atmospheric systems given in Tg N [47]

Terrestrial: Plant biomass 1.1-1.4 . 104

Animal biomass 2.102

Litter 1.9-3.3. 103

Soil: organic matter 3.105

insoluble inorganic 1.6· 104

microorganisms 5.102

Rocks 1.9.1011

Sediments 4.108

Coal deposits 1.2.105

Oceanic: Plant biomass 3.102

Animal biomass 1.7.102

Dead organic matter dissolved 5.3.105

particula te 0.3-2.4 . 104

N2 (dissolved) 2.2.107

N20 2.102

N03 5.7.105

NOZ- 5.102

NHt 7.103

Atmospheric: N2 3.9.109

N2 0 1.3 . 103

NH3 0.9 NHt 1.8 NOx 1-4 N03 0.5 Org.-N 1

Atmospheric Inventories

69

The atmosphere is divided into many parts, the lower two being the troposphere and the stratosphere. In the troposphere we find approximately 75% of the total mass of the atmosphere. This part has a vertical extension of approximately 8-15 km. The stratosphere is the part above the troposphere up to approximately 50 km. In this part of the atmosphere, important photochemical processes take place and it is here that the major part of the ozone shield of the earth is found. The troposphere has rather good vertical mixing, a typical time for ground-to-top mixing of the troposphere being little more than a month. The horizontal mixing outside the subtropical belts is of the order of 100 days. The mixing between the two hemispheres is of the order of two years. These characteristic mixing times should be kept in mind when one speaks of the spatial variability of compounds found in the atmosphere. Ofthe combined nitrogen found in the atmosphere, only 0.4% is in other froms than nitrous oxide (Fig.4). These compounds are mainly

70

>99.99%

N2

<0.01% combined

R. Söderlund and T. Rosswall

5.2

1,7%

gaseous

37%

particulilte

16% org;lnic

i norganic

Fig.4. Partitioning of the various forms of nitrogen in the atmosphere. Units given are in Tg (10 '2 g) of N

ammonia, 0.9 Tg, and NOx, 2 Tg, in gaseous forms; and ammonium; 1.8 Tg, and nitrate, 0.5 Tg, as particulates. Amines are among the organic compounds which have been identified. Although present in small amounts, all these compounds play an important role in the chemistry of the lower atmosphere [30].

The Aquatic System

The oceans cover roughly 70% of the earth's surface and have a slow turnover. A water moleeule spends on the average approximately 20 years in the surface waters, which extend to a couple ofhundred metres' depth; while once incorporated in the deep waters, it takes hundreds to thousands ofyears before it leaves this compartme nt. Light penetrates only the top 100-200 m ofthe water body, and hence most of the biological activity is found in these parts of the oceans. Biological activity not utilizing light occurs in other parts ofthe water, for example, the denitrification observed in the North Atlantic is believed to take place at a couple ofhundred metres' depth [19].

The oceanic inventories of nitrogen are given in Fig.5. The pool of nitrogen in dead organic matter is large while the biomass pool is

smalI, which indicates a rapid turnover ofthe order of days to weeks for living matter in the seas.

The inorganic nitrogen is mainly nitrate, which in deeper water is not utilized to any greater extent. This nitrogen is returned to the photo-active zone in the upwelling areas of the oceans.

The Nitrogen Cycles

2.3.10 7

95.2% N2

4.8%

combined

1.1,106

47.9%

organic

52.1%

inorganic

5.3.105

99.9% dead organic matter

44% plants

56% animals

71

Fig. 5. Partitioning ofthe global inventories of nitrogen in the aquatic system. Units in Tg of nitrogen

Ammonium nitrogen is rarely found in coastal areas, even though this is the most common form of nitrogen in the excreta of aquatic organisms [11].

The Terrestrial System

The terrestrial system is highly complex and the structure of the different biogeographical zones with regard to nitrogen distribution varies greatly. The distribution of nitrogen is, of course, very different in aboreal forest as compared with a desert system or a tropical rain forest. In a model of the global nitrogen cycle, one has to disregard these inhomogeneities. This has been kept in mind as a limit of the model when interpreting the data.

The distribution of nitrogen in various compartments of ecosystems varies according to the bioclimatic regions as given in Table 4. It is characteristic that proportionally more nitrogen is found in plant biomass as compared with the aboveground litter pool when the mean annual temperature is increased. Thus, the nitrogen pool in litter is nine times that of biomass in polar areas, while it is only 4% of that in biomass in tropical regions. H umidity also influences the total amounts of nitrogen found in an ecosystem, and the total amount of nitrogen decreases in arid conditions.

The rates of nitrogen transformations in different areas can be seen from the estimates ofyearly plant root uptake (Table 4), which is 18 times higher in tropical regions as compared with polar. This is, however, also highly dependent on the moisture regime, and nitrogen uptake in arid tropical regions is only twice that of polar regions. The largest differences between the bioclimatic regions are shown


Table 4. Distribution of nitrogen (g m - 2) between plant biomass and above-ground litter and plant uptake in different bioclimatic zones. Calculated from Bazilevich [2] and Söderlund and Svensson [47]

Biomass (gNm- Z)

Polar areas 12 Boreal areas 15 Sub-boreal areas 57

humid 137 semi-arid 22 arid 13

Subtropical areas 73 humid 161 semi-arid 68 arid 22

Tropical areas 165 humid 287 semid-arid 88 arid 8

Total terrestrial 94

3.4·10 5

93.5% organic

6.5%

inorganic

Litter Turnover time Plant uptake (g N m- 2) in li tter (yr) (g m - Z yr- 1)

106 66 1.6 76 12 6.4 10 1.0 10.0

22 14.5 5 11.0 3 4.3

9 0.4 21.2 18 37.8 10 20.7 3 11.0

6 0.2 29.3 9 46.4 6 21.6

3.4 23 1.3 17.2

96% dead organic matter

1.3.104

/ 94% ptants

4% biomass ~~~~:-:2 Va

microorganisms animals

Fig. 6. Partitioning ofthe global inventories of nitrogen in the terrestrial system. Units in Tg of nitrogen

by the values for turnover times of nitrogen in plant litter (N in litter divided by annual uptake of nitrogen, yr), the value being 66 yr in the polar regions but only 0.2 yr in the tropics.

In summary, the distribution between the various compartments in the terrestrial system is given in Fig. 6. The largest pool is the organic nitrogen, only a small part of which is biomass. The pool of nitrogen in microorganisms is very small, but this part is most important in determining the rates of transfers in the biosphere.


Only 500 Tg of nitrogen occurs in terrestrial microorganisms, but this population regulates the majority of transfer in the biogeochemical cycles, with an annual turnover of nitrogen of some 5,800 Tg [39].

Global Fluxes

In describing the nitrogen cycle as consisting of exchanges between the three reservoirs, the focus is put on the small transfer between these systems rather than on the large internal circulation within them. The fluxes within the systems are many times larger than the exchange rates. The circulation of nitrogen within the terrestrial system is at least 10 times larger than its exchange with the aquatic and atmospheric environments. The nitrogen cycle can be treated as four subcycles with different classes of nitrogen compounds:

i) ammonia/ammonium compounds, ii) NO. compounds - compounds related to NO/N02 ,

iii) N 2/N20, and iv) organic nitrogen compounds. The rationale for this subdivision is based on the small rates of interchange be

tween these classes of compounds within the atmosphere.

The Ammonia Cycle

The main source of atmospheric ammonia is the decomposition of organic matter by microorganisms, and animal excreta. Other sources are emissions from coal combustion and losses from the manufacture and use of fertilizers.

The sinks for atmospheric ammonia are wet and dry deposition. Wet deposition is the process whereby matter is transferred from the atmosphere to the surface by falling hydrometeors; dry deposition is the transfer by non-hydrometeors, such as the gaseous uptake by plants, the deposition of particulate matter by gravita-

3-8~ -r Q '",<"-".

1n- 244 I "".,. ,,~ 61-126 Y dep. I

8- 25

Terrestrial '\j2~_~1 _~~~ Fig.7. Ammonia/ammonium fluxes in Tg N yr- 1. (Söderlund and Svensson [47]). The range for the flow from the terrestrial system was obtained by balancing the other flows through the atmosphere


tional settling, and inertial impaction on to surfaces. A minor transfer of ammonia to NOx-compounds takes place in the troposphere via photochemical processes.

The exchange of gaseous ammonia between plant leaves and the surrounding air has been demonstrated by Farquhar et al. [21] to be a reversible process. At high ambient concentrations, a net transport of ammonia from the atmosphere into the plant takes place, while at low concentrations the transport proceeds in the opposite direction. The gas-phase concentration in the absence of net transport is called the compensation point. The compensation point for the species so far investigated - maize and green peas - is of the order of a few ppb(v).

A summary of the global flows of ammonia compounds is given in Fig.7.

The NOx Cycle

The principal sources for nitrogen oxides in the atmosphere are emissions from soils, the formation of NO and N02 during combustion, and lightning in the atmosphere. The sinks from the atmosphere are dry and wet deposition processes.

In field experiments, the loss of fixed nitrogen from soils, mainly nitric oxide, was demonstrated by Galbally and Roy [23]. The mechanism suggested for this loss is that nitrite in soils self-decomposes, giving off various gaseous nitrogen compounds, including NO, which have a low solubility and thus enter the intrasoil atmosphere. These gases are transferred into the atmosphere by diffusive processes. Nitric oxide can also be formed during denitrification.

The largest anthropogenie source of NOx is the combustion of fossil fuels, the global source strength of which is 19 Tg N annually (data for 1970).

The sources are unevenly distributed over the globe and 85% ofthese emissions occur in between the latitude belts of 30-60° N [37]. The global annual mean rate

Tropopause

dry dep.

13-30 19-53 0-79 + 19

Terrestrial

0.3

• f ~ 10-40

dry dep.

5-16

6-17

t~ Fig.8. Simplified flow diagram for NO, eompounds in Tg N yr- 1. Revised from Söderlund and Svensson [47]. The input due to lightning was taken from Chameides et al. [8]. The output range from the terrestrial part was obtained as a residual souree for the flows through the atmosphere and the direet anthropogenie eontribution added as indicated ( + )


for this flux is 0.15gNm- 2 yr-1, which shou1d be compared to the va1ue of 1.5 g N m- 2 yr- 1, found by Söderlund [46] from amass-ba1ance study ofNW Europe.

NOx emission rates from the burning of forests is 1ess known. Vines [54] gave data on the chemical composition of flue-gases from bush fires in Australia. By comparing the ratio of nitrogen oxides to excess carbon dioxide he concluded the emissions of the NOx gases to be less than 1/600 of the carbon source.

By using the given ratio and the estimate for the global emission of CO2 from the burning ofvegetation as given by Woodwell et al. [55] of 5,000-10,000 Tg, an upper estimate of 17 Tg N for the global emissions ofNOx from this type of source is obtained.

The fixation by lightning on a global basis has been estimated at 10-40 Tg N yr- 1 [8].

Figure 8 summarizes the global NOx cycle.

The principal processes that involve these compounds are biological and industrial nitrogen-fixation, denitrification - both biological and chemical- and to some extent nitrification. The industrial fixation is described in an earlier section, and the present rate of fixation is estimated at 45 Tg N [27]. The biological nitrogen fixation is estimated at 139 Tg N (Table 5). The increased use of legumes in agricultural practices has probably increased this flux by at least a factor of two as compared with a global system without agriculture.

The end-product of denitrification is either nitrous oxide or dinitrogen. In the atmosphere, the turnover time for dinitrogen is millions of years, and for nitrous oxide tens to hundreds ofyears. Actual measurements of denitrification in situ have not been made to any large extent, and the rates given for this flux in the nitrogen cycle have been made by using mass-balance studies that assume steady-state conditions. Thus, the estimates are very uncertain. Recent estimates for global denitrification ranges between 100 and 575 Tg N yr- 1 (Table 5).

The emissions of nitrous oxide from different systems have been measured. Hutchinson and Mosier report [28] on rates of 0.26 g N 20-N m - 2 of crop from

Table 5. Global nitrogen fixation rates for terrestrial systems [5]

System

Agriculture: legumes rice other crops grasslands

Forest Others Total

Rate (Tg yr- 1)

35 4 5

45 40 10

139


a fertilized cornfield (application rate of 20 g N m - 2), using two independent measuring techniques. They found a high variability of N 20 emission during the period studied. About 50% of the emissions were found during a one-week period directly after the application of the ammonia fertilizer. This could possibly be attributed to losses during nitrification, as pointed out by Bremner and Blackmer [4]. A second peak was observed directly after irrigation. Scaling their mean rate of N20 emission observed to aIl agriculturalland areas, and assuming the rate for noncultivated land areas to be one-fourth of that, the same authors came to a figure of 4 Tg N of nitrous oxide emissions from agriculturalland areas and 16 Tg for the global terrestrial areas. From oceans, Cohen and Gordon [9] estimated N 20 emissions from nitrification to be 4-10 Tg yr- I .

To check the annual total production of nitrous oxide released into the atmosphere, the empirical relation between variability and turnover of a compound in the atmosphere can be used as proposed by Junge [29]. Pierotti and Rasmussen [35] reported a low variability, which implies a long lifetime of atmospheric nitrous oxide, and they concluded that the lifetime is greater than 32 years. From this result, together with the total atmospheric inventory, an annual turnover of less than 41 Tg N20-N, is obtained, which is consistent with the figures given above.

To what extent denitrification in the oceans contributes to the nitrous oxide found in the atmosphere is not weIl known. Hahn [25], in measuring the nitrous oxide concentration in the surface waters of the Atlantic, found a supersaturation with regard to the ambient concentration, indicating an upward flux. From other ocean areas only a few measurements have been reported, giving an inconsistent picture.

One sink for nitrous oxide from the troposphere is diffusion into the stratosphere, where the major part forms dinitrogen and a small fraction gives NOx • The NOx compounds formed are returned to the lower layers of the atmosphere at a low rate, where they are deposited by wet and dry deposition processes. A more

N20

+ Tropopause 10 ---------------------------------------I "'''''' ", .... " , =r I

139 36 19 16-6991-92*

N2 (pelagic) N2 (sed.) I "j' 20-120 5-99- 20-80

j Terrestrial

Aquatic~

Fig.9. Global flows of N2/N2 0 (Tg N yr- 1). The value for flows of N2 to the atmosphere (*) were balanced assuming the system to be in a steady state (Söderlund and Svensson [47])

The Nitrogen Cycles

Table 6. Organic nitrogen compounds in aerosols in the 0.1-1 ~m dia. range, as detected by highresolution mass spectroscopy concentrations in ~gm-3, Cronn [10]

Alkyl nitrates Alkyl nitriles Amides Alkyl piperidines

quinoline 2-methyl imidazoline r-substituted piperidines 1-4 benzoquinone

o -0.07 o -0.3 0.005-1.5 0.01 -1.12 o -0.06 o -0.05 o -0.04 o -0.1

77

complete picture of the transport and photochemical reactions in the atmosphere is given in separate chapters in this volume.

A summary of the Nz/NzO cycle is given in Fig.9.

Organie Nitrogen Transfers

A large pool of organic nitrogen exists in both the aquatic and terrestrial systems. The transfers through the atmosphere of this type of nitrogen compound are poorly known, as weIl as their chemical composition. However, some measurements have been made of the total organic nitrogen in precipitation. The values reported are sometimes of the same order of magnitude as the total sum of inorganic nitrogen forms, pointing to the importance of this transfer of nitrogen in mass-balance studies.

Organic nitrogen compounds in natural precipitation were identified by Semenov et al. [44]. They found amino acids, amides and proteins at concentration levels of a few J..lg 1- 1 , both in snow and liquid precipation.

Troposphere

? dry dep.

? 10-100

Terrestrial

.. 10- 20 dry dep.

? ? ?

8-13

-~~~--'-------'--'~~~~~~~~~~ ,,~~~~~~~~~~~~~~~~~~~

.3s.~

~~~~~~~~~

~~~~~~

Fig. 10. The flows of organic nitrogen in Tg yr- 1 [47]


t 10

nitrouJ oxide

/1

nitrogen nitrogen biologicol t oXides oXides fixation

~ 5-30 30-130

. ammonla denitriflcat ion ammonra 90_190 orgamc lr ? 30-180

~ compounds 10 - 20

"" i

ig. 11. A global nitrogen cycle. Units are in Tg (10 12 g) yr - 1

Table 7. Global nitrogen budgets (Tg N yr- 1) [41]

Eriksson Robinson Delwiche Bums Söderlund (1959) [20] and (1970) [13] and and

Robbins Hardy Svensson (1970) [38] (1975) [5] (1976) [47]

Biological fixation: land 104 118 44 139 139 oceans n.d. 12 10 36 30-130

Atmospheric fixation (lightning) n.d. n.d. 8 10 ? Industrial fixation n.d. 20 30 30 36" Combustion 15 19 n.d. 20 19" Fires n.d. n.d. n.d. n.d. n.d. Biogenie NO, production n.d. 234 n.d. n.d. 21-89 Denitrification: land 65 n.d. 43 140 107-161

oceans 87 n.d. 40 70 25-179 Ammonia volatilization 99 957 n.d. 165 113-244 Dry deposition NH 3/NH.t n.d. 175 n.d. 72-151 Wet 99 796 140 38- 85 Dry deposition NOx n.d. 22 n.d. n.d. 25- 70 Wet 48 83 60 18- 46 Dry deposition Org.-N n.d. n.d. n.d. Wet 36 n.d. n.d. 10-100 River run off NH.t

}13 } } 15 < 1

NO] 21 30 5- 11 N-org 8- 13

N 2 0 sink: land n.d. 353 n.d. n.d. n.d. oceans n.d.

n.d. = No data given " Data for 1970 b Data for 1974 ' Data for 1976


In aerosol sampies colleeted in apolluted area, Cronn [10] found a large number of organic nitrogen eompounds (Table 6). The sources of these compounds are poorly known, with one exception, PAN, which is formed in smog situations by reactions between hydrocarbons and nitrogen oxides [43].

Dean [12] attributed an oeeanic origin to most of the organic nitrogen found in precipitation in a coastallocation. He claimed that the slick-forming material and bacteria found on the surface film were the main sources of the organic nitrogen. The eomposition of this film has been reported elsewhere as being mainly of organic material [1].

Assuming a value of 0.5 mg 1- 1 for the total coneentration of organie nitrogen compounds in precipitation and assuming a washout ratio as defined by Slinn et al. [45] of 105 , similar to that of small particles, gives an ambient concentration of 5 ng m - 3 of organie nitrogen. No data are available to confirm tbis estimate.

More research is needed in this field in order to attain a better understanding of this part of the nitrogen eycle. A summary of the organie nitrogen cycle is given in Fig.lO.

The Global Nitrogen Cycle

The global nitrogen eycle given in Fig. 11 is a summary of the subcyc1es presented earlier in the text. A number of global nitrogen cyc1es have been published reeently, and these are eompiled in Table 7. There are stilliarge uneertainties in the esti-

McElroy CAST Delwiche Liu et al. Hahn and Sweeney NAS Bolin et al. (1976) [7] (1977) [14] (1977) [31] Junge et al. (1979) [3] (1976) [32] (1977) [26] (1978) [51] (1978) [33]

170 149 99 200 180 100 139 140 10 1 30 40 85 15 -90 100 20-120 10 10 7 10 n.d. 0.5- 3 30 ? 40 57 b 40 40 40 35 70 c 40 40 20 18 20 n.d. 15 21 20

n.d. n.d. 50 n.d. n.d. n.d. 10-200 n.d. n.d. n.d. n.d. ? n.d. 22- 66 20- 90 243 70-100 120 140 150 90 197-390 63-245 106 70-100 40 130 165 50-125 0-120 35-330 150 n.d. 75 190 170 15' 36- 90 110-250

j~ jad } 79 jm, j~ jnd j8>~2

} 110-240

} 34 J 40-110

} n.d. }n.d.

} 20 }n.d. } 35 } 30 }40 } 30 } 18 } 25- 35

0 n.d. n.d. n.d. 50

n.d. n.d. n.d. n.d.

d Precipitation only e Net transfer from land to oceans


mates, and not until a better data base is available will it be possible to construct a more precise global nitrogen budget.

The transfers which are believed to be known best are the deposition rates from the atmosphere and the transfer in river run-off. The figures might have to be changed, most likely towards lower deposition, since the data on which these fluxes are based are mainly from the Northern hemisphere, where human activity has increased the rates of nitrogen transfer as compared with natural systems. The importance of natural sources for the ammonia and NOx compounds would thus be decreased, and man's impact on the global nitrogen cyc1e would be greater.

References

1. Baker, R.D., Zeitlin, H:. J. Geophys. Res. 77, 5076 (1972) 2. Bazilevich, N.I.: Energy flows and biogeochemical regularities ofthe main world ecosystems, Proc.

First Internat. Congr. Egology. Structure, Functioning and Management of Ecosystems, p.182, Pudoc, Wageningen 1974

3. Bolin, B.: Quart. J.R. Met. Soc. 105, 25 (1979) 4. Bremner, J.M., Blackmer, A.M.: Science 199, 295 (1978) 5. Bums, R.C., Hardy, R.W.F.: Nitrogen Fixation in Bacteria and Higher Plants, p.189, Springer

Verlag Berlin-Heidelberg-New York 1975 6. von Bülow, J.F.W., Döbereiner, J.: Proc. Nat. Acad. Sci. USA 72, 2389 (1975) 7. CAST: Effect of increased nitrogen fixation on stratospheric ozone. Council for Agricultural Sci

ence and Technology No. 53. Ames, Iowa: Iowa State University 1976 8. Chameides, W.L. et al.: J. Atmos. Sci. 34, 143 (1977) 9. Cohen, V., Gordon, L.I.: J. Geophys. Res. 84, 347 (1979)

10. Cronn, D.: Analysis of atmospheric aerosols by high resolution mass spectrometry. Ph.D. Disser-tation, University of Washington, Seattle 1975

11. Day, F.H.: The Chemical Elements in Nature, p.372, G.G. Harrap & Co., London 1963 12. Dean, G.A.: New Zealand J. Sci. 6, 208 (1963) 13. Delwiche, C.C.: Sci. Am. 223(3), 137 (1970) 14. Delwiche, C.C.: Ambio 6, 106 (1977) 15. Delwiche, C.C., Bryan, B.A.: Ann. Rev. Microbiol. 30, 241 (1976) 16. Döbereiner, J.: Zbl. Bakt. Parasitenk. H. 124,224 (1970) 17. Döbereiner, J.: Ambio 6, 174 (1977) 18. Döbereiner, J.: Influence ofenvironmental factors on the occurrence of Spirillum /ipoferum in soils

and roots. In: Environmental Role of Nitrogen-fixing BIue-green Algae and Asymbiotic Bacteria, (Granhall, U., ed.), Ecol. Bull. (Stockholm) 26, 343 (1978)

19. Dyrssen, D., Gundersen, K.: The nitrogen balance of the sea. In: Cyc1es (Söderberg, E., Wenzel, A.-K., ed.), p.137, Swedish Natural Science Research Council, Stockholm 1976

20. Eriksson, E.: Sv. Kem. Tidskr. 71, 15 (1959) 21. Farquhar, G.D., Wetselaar, R., Firth, P.M.: Science 203, 1257 (1979) 22. Focht, D.D., Verstraete, W.: Adv. Microb. Ecol. 1, 135 (1977) 23. Galbally, I.E., Roy, C.R.: Nature 275, 734 (1978) 24. Granhall, U., Lid-Torsvik, V.: Nitrogen fixation by bacteria and free-living blue-green algae in tun

dra areas. In: Fennoscandian Tundra Ecosystems, Part I, p. 305, (Wielgolaski, F.E., ed.), SpringerVerlag Berlin-Heidelberg-New York 1975

25. Hahn, J.: Tellus 26, 160 (1974) 26. Hahn, J., Junge, C.: Z. Naturforsch. 32A, 190 (1977) 27. Hauk, R.D.: Nitrogen fertilizer effects on nitrogen cyc1e processes. In: Terrestrial Nitrogen Cyc1es.

Processes, Ecosystem Strategies and Management Practices, (Clark, F.E., Rosswall, T., ed.), Ecol. Bull. (Stockholm) 33, 551-562

28. Hutchinson, G.L., Mosier, A.R.: Science 205, 1125 (1979) 29. Junge, C.E.: Tellus 26, 477 (1974)


30. Levy, H.: Photochemistry 9, 369 (1974) 31. Liu, S.c. et al.: Tellus 29, 251 (1977) 32. McElroy, M.B. et al.: Rev. Geophys. Space Sci. 14, 143 (1976) 33. NAS: Nitrates: An Environmental Assessment, USo Nat. Acad. Sci. Washington, D.C. 1978 34. Nutman, P.S.: Symbiotic Nitrogen Fixation in Plants, IBP Synthesis Series, Vol. 7, Cambridge Uni-

versity Press, Cambridge 1975 35. Pierotti, D., Rasmussen, R.A.: Nature 274, 574 (1978) 36. Ritchie, G.A.F., Nicholas, D.J.D.: Biochem. J. 126, 1181 (1972) 37. Robinson, E.: ACS Adv. Chem. Ser. 113, 1 (1971) 38. Robinson, E., Robbins, R.C.: J. Air Poil. Contr. Ass. 20, 303 (1970) 39. Rosswall, T.: The internal nitrogen cyc1e between microorganisms, vegetation and soil. In: Nitro

gen, Phosphorus and Sulphur - Global Cycles, (Svensson, B.H., Söderlund, R., eds.), SCOPE Rep.7. Ecol. Bull. (Stockholm) 22, 157 (1976)

40. Rosswall, T.: Aquatic mierobial eeology/Coneepts and trends. In: Microbiology 1980. Amer. Soe. Mierobiology, Washington, D.C. 1980

41. Rosswall, T.: The biogeoehemieal nitrogen cycle. In: Some Perspeetives ofthe Major Biogeoehemical Cycles. SCOPE Report 17. (Likens, G.E., ed.), John Wiley, Chichester-New York-BrisbaneToronto 1980

42. Rosswall, T., Granhall, U.: Ecol. Bull. (Stoekholm) 30, 209 (1980) 43. Seinfeld, J.H.: Air Pollution, MeGraw-Hill, New York 1975 44. Semenov, A.D. et al.: Dokl. Akad. Nauk USSR 173, 1185 (1967) 45. Slinn, W.G.N. et al.: Atmos. Envir. 12,2055 (1978) 46. Söderlund, R.: Ambio 6, 118 (1977) 47. Söderlund, R., Svensson, B.H.: The global nitrogen cyc1e. In: Nitrogen, Phosphorus and Sulphur

- Global Cycles, (Svensson, B.H., Söderlund, R., eds.), SCOPE Rep. 7, Eeol. Bull. (Stoekholm) 22, 23 (1976)

48. S0fensen, J.: Appl. Environ. Microbiol. 35, 301 (1978) 49. Stewart, W.D.P.: Nitrogen Fixation by Free-living Mieroorganisms, IBP Synthesis Ser., Vol. 6,

Cambridge University Press, Cambridge 1975 50. Stewart, W.D.P.: Ambio 6, 166 (1977) 51. Sweeney, R.E., Lieu, K.K., Kaplan, LR.: Oeeanie nitrogen isotopes and their uses in determining

the sourees of sedimentary nitrogen. In: Stable Isotopes in the Earth Sciences, (Robinson, B.W., ed.), DSIR Bull. 220, 9-26, New Zealand Dptmt. Sei. Ind. Res. 1978

52. Tiedje, J.M., S0rensen, J., Chang, Y.-Y.: Assimilatory and dissimilatory nitrate reduction: Perspectives and methodology for simultaneous measurement of several nitrogen eyc1e processes. In: Terrestrial Nitrogen Cyc1es. Proeesses, Eeosystem Strategies and Management Impacts, (Clark, F.E., Rosswall, T., eds.), Ecol. Bull. (Stockholm) 33, 331-342

53. Venkataraman, G.S.: The role ofblue-green algae in tropical rice cultivation. In: Nitrogen Fixation by Free-living Micro-organisms, (Stewart, W.D.P., ed.), IBP Synthesis Ser., 6, p. 207, Cambridge University Press, Cambridge 1975

54. Vines, R.G.: Characteristies and behavior ofbushfire smoke. In: Proc. Internat. Symp. Air Quality and Smoke from Urban and Forest Fires, p.IOI, USo Nat. Aead. Sci., Washington, D.C. 1976

55. Woodwell, G.M. et al.: Seienee 199, 141 (1978) 56. Zel'dovic, Y.B., Raizer, Y.P.: Physics of Shoek Waves and High-Temperature Hydrodynamies

Phenomena, Aeademie Press, New York 1966

The Carbon Cyde

A. J. B. Zehnder

Federal Institute for Water Resourees and Water Pollution Control (EA WAG) CH-8600 Dübendorf, Switzerland

Introduction

Carbon is a key element of life and the existence of the carbon cyc1e is the conditio sine qua non for every biological activity on this planet. The exceptional position of carbon among other elements is based on the ability of its atoms to link manifoldly together to form a huge variety of compounds which is essential to the development of living things.

About 0.1 % or 6 x 1010 million tons of the earth crust is carbon, most of it bound as carbonate in limestones or dolomites. Basically all this carbon has escaped the interior of the earth in form of gas primarily as carbon dioxide, carbon monoxide and methane. Most of the initial carbon dioxide was absorbed by the ocean and precipitated as calcium carbonate [74]. Methane, besides hydrogen, ammonia and water vapor, remained the main constituent of the atmosphere of the primitive earth. Radiation and electric discharges in such an atmosphere lead to· the formation of amino acids [e.g. 73, 79, 88, 96, 116, 130], nuc1eic acid bases [87], sugars [19, 74] and nuc1eosides, a combination of nuc1eic acid bases and sugars such as ribose or deoxyribose [91]. The nuc1eosides provide for the most essential functions in the economy of a living organism: the storage and transmission of its genetic information and hence the instruction for the synthesis of all cellular constituents. The phosphate esters of the nuc1eosides, the nuc1eotides, are organized in macromolecules, the nuc1eic acids. Among them are the largest molecules known: the deoxyribonuc1eic acid (DNA) the carrier of the genetic code of allliving things. Therefore, nuc1eosides represent the major gap in the synthesis of biologically important intermediates under primordial conditions. Beside fatty acids [37] also porphyrines could be synthesized chemically in a reducing "abiotic" atmosphere [43]. For more information about chemical evolution and abiogenic synthesis ofbiopolymers the monographs ofKenyon and Steinmann [56] and Calvin [21], the comprehensive reviews of Lemmon [61] and Degens [25], and the critical discussion on the so-called Oparism by Woese [129] may be consulted.

84 A.J.B. Zehnder

Onee life had evolved earbon was assimilated and respired again and the earbon eyele has been considerably accelerated. Beside carbon, the biosphere is also composed of the other "biological elements" (H, 0, N, P, S) and some metals in minor quantities; the carbon cyele can therefore not be dealt with independently of the cyeles of these other elements. The term biosphere is used here and later as the sum total of all plants and animals [76] rather than the place in which they live [117]. This definition allows a elearer description of the interrelationships between living things and their abiotic environments.

The growing interest in the global carbon cyele may be explained by the current concern about the "greenhouse effect" of the increasing CO2 concentration in the atmosphere as a result of fossil fuel burning, deforestation and a more intensive agricultural activity. During the past two decades much information about the functioning of the natural carbon cyele was collected, but we are nevertheless far from understanding the long-term dynamics in this cyele. Our knowledge of the global carbon cyele is summarized essentially in the artieles of Bolin [11], Garreis et al. [35], Keeling [51], Martin [68], Reiners [93], and in the SCOPE Report No. 13 [102]. The latter gives the most complete description of all aspects ofthe global carbon cyele. Numerous monographs and reviews cover the relationship between biosphere and carbon [123, 131, 133, 135] and the transformation of carbon within the pedosphere and the hydrosphere, the two major sites ofbiological activity [12, 13, 24, 26, 33, 75, 80, 90, 92, 99, 120, 132]. Some informations are available on the biochemistry of a forest ecosystem [64]. The role and influence of carbon dioxide on the global warming and carbon budget has been the subject of a number of artic1es and reviews [e.g. 1,3,18,27,34,40,39,52,53,57,65, 103, 134]. Mathematical models have been formulated concerning the CO2 interaction between atmosphere, oceans and land bio ta [9, 17,51, 104] and the consequences ofman's alteration of the carbon cyele are elucidated in areport of the Dahlem Konferenzen [110].

The objective of this chapter is to briefly summarize the global carbon budget and to draw attention on the biological processes of carbon dioxide production and eonsumption as well as on their biogeochemical importance.

The Global Carbon Cycle

Table 1 summarizes the carbon content and mean residence time for the major reservoirs. Figure 1 gives a schematic diagram ofthe main fluxes. Numerous estimates have already been made for the world net primary production. The newer estimates vary considerably. The highest values are reported by Bolin et al. with 108 x 1015 g C yr -1 (Land 63, sea 45) and Bazilevich et al. [7] with 105 x 10 -15 g yr- 1 (Land 78, sea 27). Most estimates are elose to 78.4 (Land 53.4, sea 25) given by Whittaker and Likens [125]. Garreis et al. [35] divide their 60 x 1015 g yr- 1 into equal parts for land and ocean production. Whittaker and Likens [123] summarized all published estimates of the world primary production. An updated table can be found in Ajtay et al. [1].

The carbon cyele can be roughly divided into two cyeles, a biological cyele and a geological cyele:

The Carbon Cycle

Table 1. World carbon budget: Reservoirs and residence times'

Car bon reserv oir

Atmosphere

Carbon dioxide Methane Carbon monoxide Unknown

Total atmosphere

Land

Living biomass (plants) Other living biomass and dead organisms

Total land

Ocean

Living biomass (plants) Particulate organic carbon (POC, assuming 20 Ilg Cll) Dissolved organic carbon (DOC, assuming 700 Ilg/l) CO2/HC03 Totalocean

Lithosphere

Carbonates Organic material b

Carbon content (1015 g Cl

648 6.24 0.23 3.4

657.87

827 1,200

2,027

17.4 30

1,000 38,400

39,447.4

60.9 X 106

12.48 x 106

• Values from GarreIs et al. [35], Williams [128] and Woodwell et al. [134]

85

Residence time (year)

4 3.6 0.1

16 40

0.07

385

b From this value fossil fuel is only a small percentage sufficiently concentrated to be extracted

Before man started to burn fossil fuel the geological cycle was balanced. An equal amount ofthe CO2 released through rock weathering was precipitated as calcium carbonate. Fossil fuel burning adds 5 x 1015 gC yr- 1 to the atmosphere (1978) or about 6.6% ofthe carbon released from the biosphere by respiration and decay. Also, fossil fuel burning accelerates the removal of organic carbon from the lithosphere; this amount, however is minuscule compared to the total reservoir (12.48 x 1021 g C).

The biological cycle is short but intense and thus very significant. Nearly all of the assimilated CO2 carbon is returned from the biosphere to the other spheres and the cycle is essentially closed. Solar radiation keeps this cycle active by providing the energy which is harnessed by photosynthesis. There, relatively simple inorganic chemieals are combined in biochemical processes to form a vast variety of organic molecules; these, in turn, feed the heterotrophie organisms. After death the constituents of living organisms are mineralized again by microorganisms. AIthough there is diversity and increased biochemical complexity as we pass from bacteria to plants and animals, their cellular metabolism and the structures concerned with the basic reactions of live have much in common. Yet, their biosynthetic and degradative powers differ markedly. For instance, only a highly speci-

86

+ +

8 +

8 <: ·2 5 ())o

_Iri CI>

'" U.

ATMOSPHERE I PEDOSPHERE I HYDROSPHERE

.!!! :g 78.4

-= <: ~ o '" ~8 CL

BIOSPHERE

8 / 8, \O~

Dead

(ä-d .;:

'" CD

LI THOSPHERE

Fig. 1. The global carbon cyc1e. Fluxes are given in 1015 g C yr- 1

A. J. B. Zehnder

fied group of microorganisms degrade the large quantities of insoluble polymers (e.g. lignins, tannins) which are uniquely made by plants.

The dynamic sequence of the biological carbon cycle has evolved over a geological time scale and in a variety of physical conditions. As examples of these conditions a terrestrial (cultivated grassland) and the oceanic cycle are discussed. The processes which govern a sedimentary cycle will be described later.

Carbon Balance in a Terrestrial Ecosystem

Table 2 gives estimates of the standing crop, the net primary production, and the residence time of the carbon in some terrestrial communities. Generally, the carbon input into an ecosystem starts with photosynthesis (Fig.2). Plants assimilate carbon during the day and respire part of it, and the latter becomes the sole process during the night. The remaining assimilated carbon (the net primary production)

Tab

le 2

. S

tand

ing

cro

p a

nd

net

pri

mar

y pr

oduc

tion

(N

PP

) of

terr

estr

ial

ecos

yste

ms a

Eco

syst

em t

ype

Are

a P

lant

mas

s b

Tot

al p

lant

N

PP

(1

01

2 m

2)

(10

3g

Cm

-2

) m

ass

(gC

m-

2y

r-l

)

(10

15

g C

)

For

ests

: T

ropi

cal

rain

17

.0

3 -3

6 34

4 45

0 -1

,600

T

ropi

cal

seas

onal

7.

5 3

-27

117

450

-1,1

25

Tem

pera

te e

verg

reen

5.

0 3

-90

79

270

-1,1

25

Tem

pera

te d

ecid

uous

7.

0 3

-27

95

270

-1,1

25

Bor

eal

12.0

3

-18

108

180

-90

0 W

ood

-an

d s

hrub

land

8.

5 1

-9

23

113

-54

0 T

empe

rate

gra

ssla

nd

9.0

0.1

-2.

3 6.

3 90

-

675

Sav

anna

15

.0

0.1

-7

27

90

-90

0 T

un

dra

an

d a

lpin

e 8.

0 0.

05-

1.4

2.3

4.5-

18

Des

ert

and

sem

ides

ert

scru

b 18

.0

0.05

-1.

8 5.

9 4.

5-11

3 E

xtre

me

dese

rt (

rock

, sa

nd,

ice)

24

.0

0 -

0.1

0.2

o -

4.5

Cul

tiva

ted

land

14

.0

0.2

-5.

4 6.

3 45

-1

,800

S

wam

p an

d m

arsh

2.

0 1.

4 -2

2.5

13.5

36

0 -2

,700

L

ake

and

str

eam

2.

0 0

-0.

05

0.02

45

-

675

Tot

al t

erre

stri

al

149

827.

5

a S

tand

ing

crop

an

d p

rodu

ctio

n va

lues

fro

m W

hitt

aker

and

Lik

ens

[123

, 12

5J,

and

Whi

ttak

er [

124J

b

The

car

bon

cont

ent

was

tak

en a

s 0.

45 o

f the

bio

mas

s ,

The

eff

icie

ncy

is g

iven

as

10

-3 g

ass

imil

ated

car

bon

per

g ca

rbon

of

the

stan

ding

cro

p an

d ye

ar

d R

esid

ence

tim

e o

f th

e ca

rbon

in

the

plan

ts o

f the

ir e

cosy

stem

Tot

al N

PP

P

hoto

synt

heti

c (1

01

5 g

C y

r-I )

ef

fici

ency

' (1

0-3

gg

-I y

r-I)

16.8

49

5.

4 46

2.

9 37

3.

8 40

4.

3 40

2.

7 11

7 2.

4 38

1 6.

1 22

6 0.

5 21

7 0.

7 11

6 0.

03

150

4.1

651

2.7

200

0.4

20,0

00

52.8

Mea

n re

side

nce

tim

e of

Cd

(yea

r)

20.4

21

.7

27.0

25

25

8.5

2.6

4.4

4.6

8.6

6.7

1.5

5.0

0.05

15.7

..., P" " n po ..., cr"

0 ::; n '-< n. " 00

-.

.I

88

Turnover in years

Turnover in 10' s of years

Respiration. CH4 Excretion

Turnover in loo's of years

Turnover in 1000's of years

Assimilation

Plants 1.8

~--UU

Soil

0.41

'" I I I I I I

A. J. B. Zehn der

Respiration

0.004

'" I I I I I I

, 0.005

'"

I I I I I

/' / I / I

/ I / I

0.003! I I I

Fig. 2. Carbon dynamics for the 0-20 cm layer of a grassland soil. Pools are given in kg C m - 2 and the fluxes in kg C m - 2 yr- 1. Modified from Schlesinger [99]

is used to build plant material. Parts ofthe plants are eaten by ruminants and other animals and subsequently respired or excreted. The remaining part of the plant enters into the soil. The bulk of the organic matter in most soils consists of humus or humic substances. These are amorphous, brown or black, hydrophilic acidic, polydisperse substances of molecular weights ranging from several hundreds to tens of thousands .. Humic substances are secondary products of microbial activities, especially fungi. The humification process might take place as folIows: Lignin is degraded by microorganisms to phenolic substances. They undergo enzymatic or extracellular autoxidation reactions to form highly reactive radicals or hydroxybenzoquinones which condensate with other phenolic units, peptides and amino acids to form the large humic acid molecules. The possibilities for reactions or combinations are almost unlimited. Every molecule of humus can be different but has similar properties related to the numerous active groups such as carboxyl, phenol and hydroxyl. The formation ofhumus depends on factors such as physical and chemical properties ofthe soil, vegetation, population and activity ofmicroorganisms, and on the hydrothermal conditions. Based on their solubility in alkali and acid, humic substances are usually divided into three main fractions [101]: (i) humic acid, which is soluble in alkaline solution but precipitates after acidification of the alkaline extract; (ii) fulvic acid, which remains in acidic aqueous solutions and (iii) humin which is insoluble in either acid or base.

The Carbon Cycle 89

The distinction between humic and fulvic acids, however, depends on many chemical and physical factors e.g. concentrations of the humus extraction, ion activity in the extract, temperature ofthe extract, etc. Therefore, the definition ofthese two acids which derives only from their separation procedure should just be taken as an operational distinction [22].

The humic substances are preferentially degraded by certain fungi, however this process is very slow. The mean residence time for the "mobile" humus fraction (humic and fulvic acid) is between 250 and 800 years [2, 69], whereas the unextractable humins are often older than 2,000 years [2].

The most recent estimates for the quantities of organic carbon accumulated in soil (1 m depth) are 2,946 ± 500 x 1015 gC (Bohn [10]) or 2,070 x 1015 gC (Ajtay [1]). The actual concentration of carbon in different ecosystems varies considerably. Desert (semidesert) and equatorial rain forest have a mean carbon content of 8,000 gern - 2 measured to 1 m depth. F or swamps and marshes as weIl as for temperate dry grassland values of 30,000 gern - 3 are reported [1].

The weight of microorganisms compared to the total organic carbon content of the soil is very small. Duvigneaud and Bowen [28] give aglobai estimate of 3.4 x 1015 gC and Ajtay et al. [1] base their calculation on Rosswall [97] and find 6.6 x 1015 gC bound in microbes. According to FitzPatrick [31] there are 200-1,200 g m - 2 microorganisms (live weight) in the first 15 cm of the top soil. This amounts to 10-60 gern - 2 which is about equally distributed between bacteria and fungi.

Carbon Balance in the Ocean

The ocean should be considered as consisting of two layers: an upper, euphotic zone and the deep ocean waters which are below the thermocline. These two layers show considerable differences in kind and rates ofthe carbon turnover. The surface waters are in direct contact with the atmosphere and the physical conditions are quite favorable for biological activities. The abyssal waters, however, represent a more or less isolated system. Its exchange rate with the overlaying waters is extremely slow and the rates of biologically mediated processes are considerably smaller due to the high hydrostatic pressures [49], the absence oflight, and the low temperature.

The driving force for the marine carbon cycle is the same as for the terrestrial, namely photosynthesis (Fig. 3). Most of the estimates on marine net primary production lay between 22 and 28 x 1015 g carbon per year [1, 123]. Whittaker and Likens [125] and Woodwell [134] base their world carbon budget calculations on 25 x 1015 gC yr- 1 . Platt and Subba Rao [86] give an estimate of31 x 1015 gC yr- 1 .

Fogg [32] claims that these values might be too small since not all authors appear to have taken extracellular products into account, but he omits to give a correction factor. Assuming that 100 g carbon is fixed per m2 and year, Williams [128] calculates a net primary productivity of 36 x 1015 gC yr- 1 . Bolin et al. [12] adopt an average of 45 x 1015 gC yr- 1 , however, without giving a precise base for their estimate. About 88% of the fixed carbon are recycled in the biosphere with a mean residence time of about 26 days (Tables 1 and 3). This is considerably less

90 A. J. B. Zehnder

Evatlon COz,CO,etc

Absorption CO2 ATMOSPHERE

Rivers 2,2°k, -433% ~435"1o

HC03 ,DOC, POC

Carbon Pool of the Euphotic Zone (essentiall, C02 and HC03 )

"-Respiration o u ~ ~

8' 8

-c:: o

N

S c:: o ..,

EUPHOTIC ZONE (O-200m)

:~ 2-4°/' o ° c..

:~ 2-4"10 S 06%

o '" E o '-'

---~

E U o <> c::

----------~ ----,g '0 c::

'"

~ 'e.. '0 Cl>

ci::

THERMOCLINE

DEEP OCEAN

Fig, 3. Fate of the carbon in the ocean. The fluxes are given in percent of the net primary production of the euphotic zone

than the 15.7 years of the terrestrial biosphere. In terms of radiant energy reaching the earth's surface, the seas have a lower efficiency in primary production namely about 0.2% as compared to 0.4% for the land [32].

The organie carbon which reaches the deeper parts of the ocean can roughly be divided into partieulate organic carbon (POC) and dissolved organie carbon (DOC). The conventional definition of POC and DOC is based on filtration through a 0.45 11m filter. The natur and the chemistry of DOC and POC have been extensively discussed in reviews of Degens and Mopper [24], Mopper and Degens [75], Parsons [82], Riley [95] and Williams [128]. Their findings can be summarized as folIows:

DOC: Either it is actively excreted by the plankton or it enters the water through lysis of dead cells. Most of the DOC is further utilized by heterotrophie

Tab

le 3

. S

tand

ing

crop

an

d n

et p

rim

ary

prod

ucti

on (

NP

P)

of m

arin

e ec

osys

tem

s a

Eco

syst

em t

ype

Are

a P

lant

mas

s b

Tot

al p

lant

N

PP

(1

01

2 m

2)

(10

3 g

C m

-2)

mas

s (g

C m

-2

yr-

I)

(10

15

g C

)

Ope

n oc

ean

332.

0 0

--D

.002

3 0.

45

0.9

-18

0 U

pwel

ling

zon

es

0.4

0.00

23--

0.04

5 0.

0036

18

0 -

450

Con

tine

ntal

she

lf

26.6

0.

0005

--0.

018

0.12

90

-

270

Aig

al b

eds

and

ree

fs

0.6

0.01

8 -1

.8

0.54

22

5 -1

,800

E

stua

ries

1.

4 0.

005

-1.8

0.

63

90

-1,8

00

Tot

al m

arin

e 36

1 1.

74

a S

tand

ing

crop

an

d p

rodu

ctio

n va

lues

fro

m W

hitt

aker

and

Lik

ens

[123

, 12

5] a

nd

Whi

ttak

er [

124]

b

The

car

bon

cont

ent

was

tak

en a

s 0.

45 o

f the

bio

mas

s ,

The

eff

icie

ncy

is g

iven

as

10

-3 g

assi

mil

ated

car

bon

per

g ca

rbon

of t

he s

tand

ing

crop

and

yea

r d

Res

iden

ce t

ime

of th

e ca

rbon

in

the

livin

g pa

rtic

ulat

e or

gani

c ca

rbon

of t

he e

cosy

stem

Tot

al N

PP

P

hoto

synt

heti

c (1

01

5 g

C y

r-I )

ef

fici

ency

' (1

0-3 g

g-I

yr-

I)

18.7

41

,555

0.

09

25,0

00

4.32

36

,000

0.

72

1,33

3 0.

95

1,50

8

24.7

8

Mea

n re

side

nce

tim

e of

Cd

(yea

r)

0.02

4 0.

04

0.02

8 0.

75

0.66

~0.07

;l " ("l po

3- o ::;

("l

'(l <> :::

92 A. J. B. Zehnder

microorganisms. A small fraction is converted into high-molecular weight complexes as a sort of marine humus [77]. In the deep waters the degree of inertness is increased by biotic and abiotic oxidation processes in combination with condensation reactions (see also humus formation). Thus, the molecular weight increases with depth. The weight distribution shows above 5,000 m a peak at 1,500 and at greater depth one at 10,000. The average "14C_age" of deep water DOC ('" 2000 m) is 3,400 years [127] or even as old as 4,000-6,000 years [75].

POC is composed of a labile rapidly recycled fraction (living and dead plankton, and faecal pellets) and a refractory fraction. The latter can be formed in part from DOC by biological processes [82] or by photochemical reactions at the air-sea interface [l36]. For the wh oIe water column the biological oxidation ofPOC is estimated to 2-4 g C m - 2 yr -1 [59]. The deep sea sediments are the ultimate resting place for the remaining refractory POC in the ocean. Menzel [72] estimated that the annual quantitiy of organic carbon which reaches the sediments as particulates is approximately 0.6 x 1015 g C yr- 1 . This results in a global average carbon flux to the sediments of 1.7 g m- 2 yr- 1 .

Carbonate Deposition

Organisms (e.g. Coccolithophores) are responsible for nearly all the calcium carbonate precipitation in the ocean. Most surface waters in the oceans are 4 to 7 times supersaturated with respect to calcite and argonite [63], whereas deep ocean water is normally undersaturated with respect to these minerals, largely due to the effect of decreasing temperature, increasing pressure [47] and a lower pH caused by progressive oxidation of organic matter [81]. Roughly 80% ofthe calcium carbonate formed in the surface waters of the oceans dissolve in the deeper parts [38]. The remaining 20% is probably transported by means offaecal pellets [45]. Grazers on coccoliths excrete the calcareous remains in packages which sink much faster than discrete coccoliths [12]. Since the dissolution process of calcium carbonate is very slow in the deep ocean waters, the pellets with the cocco1ith rests re ach the ocean floor without being dissolved.

Photosynthesis

Quantitatively and qualitatively the most important biochemica1 process on earth is photosynthesis. Photosynthetic reactions capture energy from solar radiation and store it in organic molecules which are synthesized from carbon dioxide and water. Photosynthesis can be regarded a local and time-limited reversal in the universal drift towards chaos, because it produces negentropy (lack of entropy) through the formation of greater molecular orderliness. Only this steady shift away from the chemical equilibrium allowed the biosphere to develop and reach the complexity and variety of present natural systems. The photosynthetic productivity of organic materials is enormous: One m2 of green leaves synthesizes one gram sugar per hour, which is on a yearly and global base a net increase in organic carbon of 15 times the actual fossil fuel mining.

The Carbon Cycle 93

Photosynthetic Energy Conversion

Solar radiation reaching the earth surface has a wavelength between 290 nm and 4,000 nm. At sea level, about 75% of the total energy of sunlight is contained in light ofwavelengths between 400 and 1,000 nm (the visible and near infrared portions ofthe spectrum) and it is within these limits that the pigments responsible for light capture in photosynthesis have their effective absorption bands (Table 4). All photosynthetic organisms contain pigments belonging to at least two different chemical dasses: Chlorophyll and carotenoids are universal components of the photosynthetic pigment systems; some algae are supplemented by a third dass, the phycobiliproteins. Chlorophyll plays a dual role in photosynthetic energy conversi on namely as a light-gathering pigment and as a site of initial photochemical event. Carotenoids and phycobiliproteins, however, function uniquely as lightgathering pigments, passing the light energy which they absorb to the reaction center, the chlorophyll. The photosynthetic conversion of carbon and hydrogen oxides, with their low free energies, into a metastable but energy rieh carbohydrate follows this apparently simple equation:

(1)

However, the biochemical process of photosynthesis is quite complicated and many intermediate steps in that reaction are still unknown. Basically, photosynthesis can be devided into three individual processes:

The Cyclic Photophosphorylation

The pigments of the system I (Fig.4) absorb light and transfer the energy to the active center of the photosynthetic unit (P 890 of the bacteriochlorophyll a or P 700 of the chlorophyll a). The chlorophyll molecule is oxidized by ejection of an electron to a higher energy level (-0.4 V). There it reduces ferredoxin, a low potential electron acceptor. The re oxidation ofthe ferredoxin releases approximately as much energy as the oxidation of molecular hydrogen. This energy can be harnessed through the intermediacy of other electron carriers (RedjOx in Fig.4) in order to generate biologically useful forms of chemical bond energy (ATP = adenosine triphosphate):

ADP + Phosphate ~ ATP + HzO

ßGo (pH 7)= -31.8 kJjmol.

(2)

If the electrons are used to reduce COz to sugars (CO z assimilation in Fig.4) the chlorophyll has to be reduced again.

Phototrophic bacteria (except blue-green algae or cyanobacteria) use for this purpose sulfide, thiosulfide, molecular hydrogen or organic compounds depending to whieh group they belong (Table 5). Cyanobacteria and green plants, however, are able to use water. Since the redox potential of the system I is not high enough to oxidize water a second system called system II is needed.

Tab

le 4

. Chl

orop

hyll

abs

orpt

ion

spec

tra

and

phot

oact

ions

'

Wav

elen

gth

Chl

orop

hyll

typ

e A

bsor

ptio

n m

axim

a O

rgan

ism

s (n

m)

in l

ivin

g ce

lls (

nm)

<

300

300-

400

Bac

teri

ochl

orop

hyll

a

360,

390

} P

urpi

e su

lfur

and

pur

pie

non

sulf

ur b

acte

ria

Bac

teri

ochl

orop

hyll

b

400

400-

500

All

type

s of

chl

orop

hyll

40

0-50

0 A

ll ph

otos

ynth

etic

org

anis

ms

500-

600

All

proc

esse

s re

lativ

ely

inac

tive

600-

700

Bac

teri

ochl

orop

hyll

a

600

} B

acte

rioc

hlor

ophy

ll b

60

5 C

hlor

ophy

ll a

67

5

700-

800

Bac

teri

ochl

orop

hyll

c

705-

74

0}

B

acte

rioc

hlor

ophy

ll d

74

5-75

5

800-

900

Bac

teri

ochl

orop

hyll

a

800-

89

0}

Bac

teri

ochl

orop

hyll

b

835-

850

900-

1,00

0

> 1

,000

B

acte

rioc

hlor

ophy

ll b

1,

020-

1,04

0

• In

par

ts f

rom

Pfe

nnig

[84

], P

fenn

ig a

nd T

rüpe

r [8

5] a

nd S

ybes

ma

[111

] b

Aft

er W

ithr

ow [

126]

Pur

pie

sulf

ur a

nd p

urpi

e no

n su

lfur

ba

cter

ia

Cya

noba

cter

ia, e

ucar

yoti

c al

gae,

pla

nts

Gre

en s

ulfu

r ba

cter

ia

Pur

pie

sulf

ur a

nd p

urpl

e no

n su

lfur

bac

teri

a

Pur

pie

non

sulf

ur b

acte

ria

Oth

er p

hoto

eff

ects

b

Des

truc

tion

of t

issu

e P

hoto

trop

ism

(370

nm

)

Pho

totr

opis

m (

445+

475

nm)

Pho

tom

orph

ogen

ic i

nduc

tion

(6

60 n

m),

Pho

tope

riod

ic a

ctiv

ity

Pho

tom

orph

ogen

ic r

ever

sal

(730

nm

)

? Hea

ting

eff

ect

'C. ?>

!-< ~

N f

E0

1 [v

olt

J

-0.4

o 2e

- ,. --r

+

0.4

r IPiQ

men

ts I

AD

P

+ 0

.81

-

ATP

SY

ST

EM

CY

CLI

C

PH

OT

OP

HO

S

PH

OR

YLA

TIO

N I

+P

2e

- 2e

-.

2H

20

20

H1

2H

+

SY

STE

M

]I

NO

NC

YC

LIC

P

HO

TO

PH

OS

P

HO

RY

LATI

ON

O

R

PH

OT

OLY

SIS

O

F W

ATE

R

H2

C-

0-®

I C-O

H

11 6

C-O

H

I

HC

-OH

I

H2

C-

0-®

6

C0

2-1

+H

20

H2

C-

O-®

O~

I

HO,C-~-OH

C=O

I

HC

-OH

I

6 C

s

H2

C-

O-®

~ CO

OH

I

12

HC

-OH

I

H2

C-

O-®

[HJ~

+24H

[

J 1

0C

3

12 A

DP

+P

-1

2H

j2

12 A

TP

H,C~O

12

HG

-OH

--2

C3

I

H2

C-

0-®

C02

A

SS

IMIL

AT

ION

HE

XO

SE

-l,6

-D

IPH

OS

PH

AT

E

H2C

-O

-®

I H

C -

---,

HO

-C-H

t

I 0

--

H-~-OH

I H

C--

---.

.J

I H

:zC

-O-®

Fig.

4.

Sch

eme

for

the

ligh

t re

acti

ons

in p

hoto

synt

hesi

s an

d th

e C

O2

assi

mil

atio

n ov

er t

he "

Cal

vin-

Bas

sham

cyc

le".

In

gre

en p

lant

s an

d al

gae

the

phot

o sy

stem

ha

s tw

o "s

hort

-wav

elen

gth"

ligh

t re

acti

on c

ente

rs.

The

chl

orop

hyll

a o

f sy

stem

II

abso

rbs

belo

w 6

80 n

m a

nd i

n sy

stem

I a

t 70

0 nm

. In

pho

tosy

nthe

tic

bact

eria

th

e pi

gmen

ts (

bact

erio

chlo

roph

yll

a) i

n sy

stem

I a

re m

ost

acti

ve a

t th

e "l

ong-

wav

elen

gth"

of

890

nm.

eS]

stan

ds f

or e

lect

ron

dono

rs o

ther

s th

an w

ater

(T

able

5).

[H]

and

[

] m

ean

the

elec

tron

an

d p

roto

n ca

rrie

r N

AD

P+

(ni

coti

nam

ide

aden

ine

dinu

cleo

tide

pho

spha

te)

in i

ts r

educ

ed a

nd

oxi

dize

d fo

rms

resp

ecti

vely

,.., i:l" '" ("J 00 8- o ::;

("J l '0

v.

Tab

le 5

. Dis

tinc

tion

am

ong

phot

osyn

thet

ic o

rgan

ism

s

Org

anis

ms

Cya

noba

cter

ia

Euc

aryo

tic

alga

e P

lant

s }

Gre

en s

ulfu

r ba

cter

ia (

Chl

orob

acte

riac

eae)

Pur

pie

sulf

ur b

acte

ria

(Thi

orho

dace

ae)

Pur

pie

no

n s

ulfu

r ba

cter

ia (

Ath

ioro

dace

ae)

• B

y so

me

stra

ins

b D

epos

itio

n ou

tsid

e th

e ce

lls

Pri

ncip

al p

hoto

synt

heti

c el

ectr

on d

onor

H20

H2S

, S20~~, H

2 (o

rg. c

ompo

unds

)'

H2S

, S20~~, H

2 (o

rg. c

ompo

unds

)'

H2

, or

g. c

ompo

unds

(S20~~)'

Pho

tosy

nthe

tic

oxyd

atio

n pr

oduc

t

O2

SO~~, S

Ob

(C0

2)

SO~~, S

Oc

(C0

2)

CO

2

(SO~~)

Rel

atio

n to

oxy

gen

Aer

obic

Ana

erob

ic

Ana

erob

ic

Tol

eran

t

c D

epos

itio

n us

uall

y in

side

the

cel

ls a

s an

int

erm

edia

te e

xcep

t E

ctot

hior

hodo

spir

a w

ith

its

sulf

ur d

epos

itio

n ou

tsid

e th

e ce

lls

Maj

or c

hlor

ophy

lls

Chl

orop

hyll

a

Bac

teri

ochl

orop

hyll

c a

nd

d

Bac

teri

ochl

orop

hyll

a a

nd

b

Bac

teri

ochl

orop

hyll

a a

nd

b

\Cl

0-,

?> ~

!=C J

The Carbon Cyc1e 97

Noncyclic Photophosphorylation or Photolysis of Water

Figure 4 shows the coupling of systems land 11. Molecular oxygen is the "waste product" of the photosystem II:

20H- ~Hzo+iOz+2e- . (3)

Electrons driven through systems 11 and I can produce energy and be used for the reduction of COz, the third process in photosynthesis:

COz Assimilation

(4)

cO z assimilation is light independent and is therefore often called "dark reaction". The role of the photo systems is to produce reducing equivalents at a low potential and the energy necessary to reduce COz to sugars. The energy produced during the photophosphorylation can of course be used also for the synthesis of other organic compounds.

Calcite Precipitation as a Result of Photosynthesis

Besides the formation of organic material the photosynthesis can also lead to a 10-cal depletion of carbon dioxide. This is especially true for lacustric environments where biological activities control the overall COz budget. Photosynthesis by green algae or cyanobacteria (blue-green algae) uses up large quantities ofCOz. Since the exchange rate of COz at the atmosphere-water interface is slow compared to COz uptake by algae or plants [29] the surface waters of an eutrophic lake can become supersaturated with respect to calcite. Megard [71] found for some Minnesota lakes that calcium carbonate saturation is probably controlled entirely by the balance between carbon dioxide assimilation and its release during respiration. With the specific ion activities of these lakes 4 mol of carbon had to be consumed to precipitate 1 mol of calcium. The condition for calcite precipitation is defined by the relation:

Caz+ + CO~- = CaC03 . (5)

At a neutral pH most of the carbonate is present in form of bicarbonate

2HC03=CO~- +COz+HzO, (6)

but the organisms use preferentially COz for carbon assimilation (Eq. 1). The combination of Eqs. (1) and (6) show how calcite can be precipitated:

(7)

The importance of CaC03 precipitation in hard-water lake metabolism has been elaborated by several workers [46, 78, 119, 121]. This phenomenon also called whiting [6] is documented for lakes [118, 122], and shallow marine environments [15].

98 A. J. B. Zehnder

Fig. 5. Whiting of Lake Michigan as a result of CaC03 precipitation. Pictures taken by Landsat-l 1973 at 16 July, 3 August, 21 August (Jeft to right). Courtesy A. E. Strong and B.J. Eadie [108]. Reproduced by permission of the American Society of Limno10gy and Oceanography

The clouding of lakes occurs mostly seasonally and can directly be related to blooms of pelagic microphytes such as green algae and cyanobacteria (blue-green algae) as was nicely shown by Strong and Eadie [108] for the Great Lakes area (Fig. 5). The role of calcite precipitation as direct result of photosynthetic activities is probably only important locally and does not markedly influence the global carbon budget. However, in large or deep basins such as the Great Lakes or the Black Sea the chalk deposits may eventually become a significant part of the geological record [55].

The Carbon Dioxide Problem

Although COz is only a trace gas in our atmosphere, it plays a key role in maintaining life on our planet in its function in photosynthesis. Since the beginning of industrialisation man has added carbon dioxide continously to the atmosphere through burning of fossil fuel. In the last three decades the concern has developed that a further increase of COz concentration in the atmosphere may have a disastrous effect on the ecology of our planet. The voluminous literature has been ex-

The Carbon Cycle 99

tensively reviewed and discussed in the 24th Brookhaven Symposium in Biology [131], the SCOPE report No. 13 [102] and by Woodwell et al. [135]. A more critical evaluation on man's alteration of the carbon and other chemical cycles has been presented at the Dahlem Konferenz [110]. The COz problem can be divided into four components: What are the sources of COz? What are the sinks for COz? How does the climate respond to a change ofthe COz concentration in the atmosphere? What are the effects of a variation of the COz content in the atmosphere on the environment?

Sources of Carbon Dioxide

In pre-industrial times the carbon dioxide cycle was essentially balanced. The carbon dioxide from respiration was fixed again by photosynthesis. Large parts ofthe oceans, such as the equatorial Pacific are outgassing COz to the atmosphere. This results from upwelled cooler waters which be ar an excess of CO z' Other parts, such as north ofthe 50 degree North in the Atlantic represent sinks ofCO z. There, surface waters from lower latitudes move northward but cool more rapidly than they can dissolve atmospheric CO2 . It is considered that this oceanic steady state has held our atmospheric COz level constant for many thousands of years [14, 50].

With the progressive civilization vast areas of forests were destroyed. Decomposition processes and woodburning converted the carbon bound in the plants into COz. The actual amount of the so released CO z is controversial. Woodwell et al. [35] estimate a net release of 7.8 x 1015 gC0z-Cyr-1 with a range of 2-18 x 1015 gCyr- 1.

Bolin [12] has based his calculations on statistics tabulated by the FAO and other agencies. He found an annual wood harvest of 0.5 to 2 x 1015 g C. In terms of CO2 input into the atmosphere this number is considerably smaller because parts of the denuded areas become covered again by vegetation and some carbon remains bound in wooden articles for decades.

5 x 1015 g C yr- 1 is the actual addition of CO2 to the atmosphere through fossil fuel burning and kilning of limestone for cement production. Since 1860 the industrial CO2 production has grown exponentially with interruptions only during the periods of the two world wars and the great depression after 1930 [3]. Cumulative 145.6 x 1015 g C have been released since 1860 (1976 value) and the atmospheric COz concentration went up in the same period of time from 290 ppm to 333 ppm [98]. This means that over 60% of the industrial CO2 output remained airborne [34]. The totallong-term increase of atmospheric CO2 has been measured by Keeling and coworkers at Mauna Loa Observatory, Hawaii and at the South Pole (Fig.6). At both places the average annual increase was about the same for a 12 year period (1959-1971); namely 3.1 % from 315.7 to 325.4 ppm at the South Pole and 3.4% from 316.1 to 326.8 ppm in Hawaii [52, 53]. The rate ofincrease today is more than 1 ppm yr- 1 [34,52,53].

Sinks of Carbon Dioxide

The oceans can both be a source and a sink for COz depending on the latitude but overall the ocean is a sink for COz (sedimentation of CaC03 and POC). It is as-

100

E e.. e..

~ Cl>

-<=

330

g. 325 o E 1ii

.S

~ 320 ';:C o :c

0:::: o -e ~ 315

'" o

335

330

u 325 ,'-, , I , ,

I

1970 YEAR ,/ ,,-'

" , " , ,

,.-" ... "

,,.-,/ '~orth Pole Mauna Loa >,,/ ./

" .. " / ,,',.' /' .... .

",,~

1970

A.J.B. Zehnder

,,'

Fig. 6. Increase ofthe atmospheric CO2 concentration at Mauna Loa, Hawaü, the South Pole [54] and the Northem Polar region [8]. Modified from BoHn et al. [12]. The insert shows the seasonal oscillation ofthe atmospheric CO2 at Mauna Loa Observatory. The low concentrations every summer are due to the removal ofC02 by photosynthesis during the vegetation period in the Northem Hemisphere. (After Woodwell [134])

sumed that the oceans absorb 1.5 X 1015 g COrC yr- 1 which is about 30% ofthe present CO2 production by fossil fuel combustion. For a long time, the terrestrial biota has been thought as a sink for 20% ofthe antropogenic CO2 , however, newer data do not agree. The existence of a new potential source of CO2 is particularly throublesome since the mixing processes in the oceans limit their capacity of CO2

absorption at least in time periods of years or decades [17, 83]. Therefore, about 20% disappears annually from the atmosphere for which the existing models cannot account for. Although, Lerman [62] suggests to consider alkaline and calcareous soils as sinks, as yet no clear concept has evolved to explain this discrepancy.

Global Warming

In the atmosphere CO2 absorbs light in the infrared region which otherwise would be reflected away from the earth, thus heating up the earth atmosphere. An increase of the atmospheric CO2 content will absorb proportionally more heat and lead to what is called the green-house effect. Already in 1938, Callender argued [22] that as a resuIt of fossil fuel burning the climate may change. From the numerous articles and models we can draw the following conclusions: If fossil fuel burning increases at a 5% rate per year as it did from 1960 to 1976 the atmospheric CO2

The Carbon Cycle

80 ., .... >-

U 0>

Ln

-g 60

c: 0

~ 40 ~

'" c: 0 <J

a; 20 ~

'" '" 0 LL

1900

4.5%

2000

E 2500 ,--,---.---.---,--,---, 0-"'--2000

{31500

2000 2200 2400 YEAR

101

Fig. 7. Industrial CO2 production expressed as utilization of fossil fuel for 4 assumed increases of the yearly consumption rates. Insert: Solid lines. Increase of atmospheric CO2 concentration as a function of the different pattern offossil fuel consumption, but without a positive feedback from the biota. After Kee1ing and Bocastow [54]. The dashed curve was obtained with the assumption that primary production increases with increasing atmospheric CO2 concentration (annually 3%) and no limitations were caused by other environmental factors such as water, nutrients, radiation or temperature [102]

content will double ( '" 600 ppm) by the year 2020 with respect to the preindustrial time of 1860 (Fig.7 [104]). As a result the average earth surface temperature will rise by about 2.5 K. The warming will especially affect the polar regions whereas at the equator the temperature will not change markedly [66,100]. In the belts of 60 to 90 0 N or S lat. the air temperature change will be approximately 2.5 times higher than the mean alteration for the Globe [18]. This is especially due to the relatively stable stratification of the lower atmosphere in these regions and the effect of the snow-temperature-albedo feedback. The reviews of Budyko and Vinnikov [18], Schneider [100] and Sekihara [103] summarize the status quo on the global warming and critically discuss the most important models.

Our knowledge on the effect ofsuch a warming are very limited, however, most scientists today agree that it may possibly change life on this planet significantly. But the opinions diverge on what and how strong the change will be, especially since the CO2 effect can partially be offset by the natural cooling cyc1es as has been observed for the period of 1940 to 1965 [66]. Further, deforestation is assumed to be in part responsible for the increase of atmospheric CO2 and hence the global warrning. But a c1eared area reflects more radiation (has a higher albedo) and therefore provokes cooling. The question as to which extent a higher nebulosity as a result of rising temperatures might affect the warrning process has not yet been convincingly answered.

102 A. J. B. Zehn der

Environmental Responses to a Variation in Atmospheric Carbon Dioxide Content

It has been thought that CO2 concentration can act as a regulator of photosynthesis [36]. With an increase of the atmospheric CO2 to 1,000-1,500 ppm - which is the CO2 saturation level of photosynthesis - net primary production could be doubled or tripled [139]. However, some adverse effects could prevent the full response of photosynthesis on the higher CO2 concentration: (i) Self-shadowing by the leaves. (ii) Photosynthesis is partially inhibited by higher oxygen concentrations (War

burg effect [115], which would at least locally result from increased primary production.

(iii) At many places photosynthesis is limited not by CO2 or light but by either micro-nutrients or water or both.

Point (iii) is contested. Garreis and Perry [36] argue that an addition of CO2

to the atmosphere williower the oceanic pH which would tend to increase the concentration ofinorganic phosphorus in the ocean [109] and thus permit phosphorus limited system to respond to increased CO2 . Further, higher evaporation is provoked by warming with the consequence of more rainfall. A favorable influence on primary production, however, depends largely on the rainfall pattern. Ifthe increase of temperature which paralleis the augmentation of the atmospheric CO2

is more than 10 °C somewhere between 20 and 35°C the net photosynthesis ofterrestrial plants could at least be doubled [39]. In addition, the temperature rise would extend the vegetation period in the temperate zones. In aquatic environments the temperature might have a more pronounced effect on phytoplankton production provided that no micro-nutrient is limiting. Konopka and Brock [58] found a doubling or even tripling of the photosynthetic activity of cyanobacteria (bIue-green algae) when the water temperature was brought from 15 to 20 oe. The additional input of organic matter into soil and ocean will increase the oxygen demand in large areas of these environments. Since the mass of oxygen is 1,700 times the mass of atmospheric CO2 , a higher primary production due to fossil fuel burning is unlikely to have a significant effect on the oxygen partial pressure [16]. As a result more areas in soil and ocean can become anaerobic. This, however, will enhance the deposition of organics and enlarge the carbon reservoir of the lithosphere at the expense of the atmosphere. A similar process might have created the coal and oil deposits during the Carboniferous Period.

Beside the biological also geological and chemical processes can to a certain extent regulate the atmospheric CO2 . At a higher CO2 level weathering of calcium and magnesium minerals is more intense. Most of the calcium from the calcium bearing silicates and sulfates is deposited as CaC03 . The weathering and deposition of calcium compounds causes thus a net CO2 demand (Table 6 [36,44]). Warming could have an adverse effect. As was pointed out by Eriksson [30] a temperature increase by 1.5 to 3 K ofthe ocean waters will reduce its solubility for CO2

and add another 44 to 87 ppm CO2 to the atmosphere. Each biological and chemical feedback mechanism could by itself prevent or enhance the rise of the atmospheric CO2 concentration. At this stage, however, we cannot predict what kind of new steady state will be reached by a simultaneous interplay of all known and unknown reactions involved in the CO2 cycle. The question whether an increase

The Carbon Cyc1e 103

Table 6. Gains and losses of atmospheric CO2 through weathering and sedimentation a

Process

Weathering Oxidation of elemental carbon Dissolution of limestones and dolomites Decomposition of Ca- and Mg-silicates

Sedimentation Deposition of elemental carbon Deposition of carbonates Deposition of Mg-silicates

Total net loss of atmospheric carbon

• After Holland [44]

Gain (10 15 gCyr- l )

0.09±0.02

0.22±0.04 0.08±0.02

Loss (1015 g C yr- I )

0.16±0.02 0.19±0.02

0.12±0.03

0.08±0.04

of CO2 in the atmosphere is overall beneficial for mankind or not remains unanswered. The opinion of Goguel [39] is worth to be considered. He thinks that the evolution went toward exhaustion of CO2 in the atmosphere which would trigger new glacial periods but the use of fossil fuel will reduce this risk.

Biological Cycle of Carbon Dioxide

With the present conditions on this planet organic matter is thermodynamically metastable. Thus, the ultimate fate of all carbon compounds is CO2 (Fig.8). The oxidation of organic compounds is catalyzed by heterotrophic organisms according to the following equation:

(8)

The electrons from this reaction are transfered to different acceptors depending on organism and environment; thereby chemical energy is released. It is harnessed by a variety of specific, biochemical pathways to support life processes. In nature these redox reactions tend to occur in order of their thermodynamic possibilities; that is, the organic material as reducing agent supplies its electrons to the lowest unoccupied electron level (Table 7). The sequential use of electron acceptors by different organisms is very common in soil, lakes, isolated bodies ofwater, and marine sediments (e. g. Kusnezow [60]).

The best studied and energetically most efficient process is the aerobic respiration, which supplies energy to animals, plants, and partially to microorganisms. Anaerobic reactions are exc1usively carried out by microorganisms (bacteria and some fungi). Anaerobes bind about 10% ofthe available energy in their biomass whereas aerobes can fix up to 50%.

The incipient reduction of oxygen is followed by reduction of nitrate. Denitrification starts only after the oxygen partial pressure has fallen below 0.5 k N rn-i (Skerman and MacRae [lOS]), and rnanganese dioxide should only be

104 A. J. B. Zehnder

o 8 I c

\ E R o

Fig. 8. The biological CO2 cyc1e

reduced after oxygen is removed. But oxygen is needed to induce the synthesis of the mangane se dioxide reducing enzym system [114] and no organism has been found which covers its energy requirement with this reaction (e in Table 7). Similar observations have been made with ferric iron [4]. In reducing environments microbial and purely chemical reduction of mangane se dioxide and ferric iron might take place simultaneously. Aseparation of the two processes is often impossible, since the reducing conditions in the environment are maintained by microorganisms. The solubilization ofmanganese and iron as a result oftheir reduction is important in the chemistry of soils and also of some interest in the formation mechanism of ferromanganese nodules. The conversion of nitrate to ammonia has a long time been thought to proceed via the synthesis of amino acids and to be important only in the nitrogen assimilation metabolism. 20 years ago, Japanese researchers have

The Carbon Cycle 105

Table 7. Redox sequence of the biologically media ted oxidation of organic matter a

Reaction - LI GO per mol Redox of electrons potential b

exchanged EO' H

(kllmol) (mV)

Aerobic respiration (A) [CHzO]'+Oz -> COz+HzO 117.5 810 Denitrification (nitrate respiration) (B) [CHzO] +4/5 NO; +4/5 W -> CO z +2/5 N2 + 7/5 HzO 112.0 750 Manganese reduction (C) [CHzO] +2 MnOz(s)+2H+ -> MnC03 +Mn2+ +2 HzO 94.5 500 Nitrate ammonification (D) [CHzO]+~ NO; +W COz +~ NH,i +~ HzO 74.0 360 I ron reduction (E) [CHzOJ+4 FeOOH(s)+6 H+ -> FeC03 + 3 Fe2+ + 6 HzO 24.3 -100 Fermentation (F) [CHzO] -> ! COz +! [CZH6 O] 23.4 -180 Sulfate reduction (sulfate respiration) (G) [CHzO] + ~ SO~ - +~ W COz+~HS- +HzO 18.0 -220 M ethane formation d

(H) [CHzO] +~ COz COZ+~CH4 16.3 -250 Proton reduction (I) [CHzO] + HzO -> COz+2Hz 3.8 -420

a Calculated for pH 7; Hz, Oz, N z• CH4 are in gaseous state at 101 kNm- z; all other substances at 1 mol kg- 1 activity

b These values give only the redox potential at equilibrium for the overall reaction. Biological reactions, however, are mostly multi-step reactions. Each step may show a different equilibrium potential, compared to the one in the Table. Moreover the "milieu interieur" of the organisms differs quite markedly from the "milieu exterieur". Therefore, organisms may already carry cut the listed reactions at much higher environmental redox potentials

, Stands for an organic compound d Reduction of COz. The decarboxylation of acetate is not an intermolecular redox reaction:

CH3COOH -> CH4 + COz

al ready pointed out that some organisms might catalyze a direct reduction and use this reaction as an energy generating system [48, 112], however, its metabolie pathway is still not known. In coastal marine sediments the activities ofthe nitrate ammonification was found to be in the same order of magnitude as those recorded for denitrification [106].

Organic compounds can also act as electron acceptors under anaerobic conditions. Pasteur was the first to realize that decomposition of organic matter in absence of oxygen and any other external electron acceptor can also be used by some bacteria to obtain energy. He called this process fermentation. In fermentation, the electron donors and acceptors are organic compounds usually genera ted from single organic substrates in the course of the intermediary metabolism. Compounds which are to be fermen ted must therefore be capable of yielding both oxidizable and reducible intermediates. Fermentation occurs as fast as the oxygen par-


tial pressure is reduced and it can be observed simultaneously with all other anaerobic degradation reactions, provided the presence of enough substrate. Thus, its place in a redox sequence as shown in Fig.8 and Table 7 is only based on theoretical calculations; in this case for alcohol fermentation.

In natural habitats sulfate reduction can be observed just after nitrate is consumed [60,94,107].

Equation H in Table 7 describes only an ecologically observable reaction, since methane is either formed by cleaving acetate or from CO2 and H2 (methanogenic substrates of minor ecological importance are formate, methanol or the methylamins) but never by a direct reduction of CO2 with organic matter [4]. To obtain methane from CO2 molecular hydrogen has to be formed first. This is done by a specific group of organisms, the so-called "obligate proton reducing acetogens" [70] or by some fermenting microbes [113]. According to Eq. I in Table 7 proton reduction is endergonic under standard conditions. This process, however, can be shifted into the exergonic range if hydrogen is removed instantaneously to keep its partial pressure low [137]. In absence of sulfate this is most efficiently done by hydrogen oxidizing methane formers. Thus, for the production of methane during anaerobic mineralization of organic compounds at least two microbial populations have to closely interact namely methane bacteria and hydrogen formers [137].

Methane has been believed to be inert under anaerobic conditions, but some recent reports suggest that sulfate might be used to oxidize methane in marine environments [5, 23, 67, 89]. These observations contradict the current theory that methane can be oxidized biologically only with molecular oxygen. Zehnder und Brock [138] presented evidence that anaerobic methane oxidation exists and that it is a type of back-reaction of methane formation. For the global carbon cycle, however, the aerobic oxidation ofmethane [42] which closes the CO2 cycle is much more important than its anaerobic counterpart.

Acknowledgement

The preparation of this chapter was aided by a grant from the Federal Institute for Water Resources and Water Pollution Control (EA WAG), Switzerland. The manuscript was typed by Barbara Stier and the figures drafted by Heidy Bolliger.

References

1. Ajtay, G.L., Ketner, P., Duvigneaud, P.: in The Global Carbon Cyc1e (Bolin, B., Degens, E.T., Kempe, S., Ketner, P., Eds.) SCOPE Rep. 13. John Wi1ey, Chichester 1979, Chap.5

2. Baer, F.E.: Chemistry of the Soil. Reinhold Pub1ishing Corp., New York 1955 3. Baes, C.F. et al.: ORNL-5194, Oak Ridge National Laboratory, Oak Ridge, Tennessee 1976,

pp. 1-72 4. Ba1ch, W.E. et al.: Microbio1 Rev. 43, 260 (1979) 5. Barnes, R.O., Goldberg, E.D.: Geo1ogy 4, 297 (1976) 6. Bathurst, R.G.: Deve10pments in Sedimento1ogy. Elsevier, Amsterdam 1971, Vol.12 7. Bazi1evich, N.!., Roding, L.Y., Rozov, N.N.: Sov. Geogr. Rev. Transl. 12,293 (1971) 8. Bischof, W.: Tellus 29, 435 (1977)

The Carbon Cyde 107

9. Björkström, A.: in The Global Carbon Cyde (Bolin, B., Degens, E.T., Kernpe, S., Ketner, P., Eds.) SCOPE Rep.l3. John Wiley, Chichester 1979, Chap.15

10. Bohn, H.L.: Soil Sei. Soe. Amer. J. 40, 468 (1976) 11. Bolin, B.: Sei. Am. 223, 125, Sept. 1970 12. Bolin, B.: Seience 196, 613 (1977) 13. Botkin, D.B., Janak, J.F., Wallis, J.R.: in Carbon and the Biosphere (Woodwell, G.M., Pecan,

E.V. Eds.) 24th Brookhaven Symp. Biology, USAEC. Conf.-720510, 1973, pp. 328-344 14. Brewer, P.G.: Oceanus 21 (No.4), 13 (1978) 15. Broecker, W.S., Takahashi, T.: J. Geophys Res. 71, 1575 (1966) 16. Broecker, W.S.: Seience 168, 1537 (1968) 17. Broecker, W.S., Li, Y.-H., Peng, T.H.: in Impingement ofMan on the Oceans (Hood, D.W., Ed.).

Wiley-Interscience, New York 1971, pp. 287-324 18. Budyko, M.l., Vinnikov, K.Ya.: in Global Chemieal Cydes and their Alterations by Man

(Stumm, W., Ed.). Dahlem Konferenzen, Berlin 1977, pp. 189-205 19. Butlerov, A.: Liebigs Ann. Chern. 120, 296 (1861) 20. Callender, G.S.: Quart. J. Roy. MeteroL Soe. 64, 223 (1938) 21. Calvin, M.: Chemieal Evolution. Oxford University Press, Oxford 1969 22. Davis, J.A.: in Contaminants and Sediments (Baker, R.A. Ed.). Ann Arbor Sei. Publishers Ann

Arbor, 1980, VoL2, pp. 279-304 23. Davis, J.B., Yarbrough, H.F.: Chern. GeoL 1, 137 (1966) 24. Degens, E.T., Mopper, K.: in Chemieal Oceanography, 2 nd ed. (Riley, J.P., Chester, R. Eds.).

Aeademic Press, London 1976, VoL6, Chap.31 25. Degens, E.T.: in The Global Carbon Cyde (Bolin, B., Degens, E.T., Kempe, S., Ketner, P., Eds.).

SCOPE Rep.l3. John Wiley, Chichester 1979, Chap.2 26. de Vooys, C.G.N.: ibid. Chap.1O 27. Duplessy, J.C., Lambert, G.: La Recherche 9, 696 (1978) 28. Duvigneaud, P., Bowen, H.J.M.: eited in [l] 29. Emerson, S.: LimnoL Oceanogr. 20, 743 (1975) 30. Eriksson, E.: J. Geophys. Res. 68, 3871 (1963) 31. FitzPatriek, E.A.: An Introduction to Soil Science. Oliver and Boyd, Edinburgh 1974

32. Fogg, G.E.: in Chemieal Oceanography 2nd ed. (Riley, J.P., Skirrow, G., Eds.). Academie Press, London 1975, VoL2, Chap.14

33. Fonselius, S.: Ambio Spee. Rep. 1,29 (1972) 34. Freyer, H.-D.: in The Global Carbon Cyde (Bolin, B., Degens, E.T., Kempe, S., Ketner, P. Eds.).

SCOPE Rep. 13. John Wiley, Chichester 1979, Chap.3. 35. Garrels, R.M., Maekenzie, F.T., Hunt, C.: Chemical Cydes and the Global Environment. William

Kaufmann, Los Altos, California 1973, Chap.6 36. GarreIs, R.M., Perry, E.A.: in The Sea (Goldberg, E.D. Ed.). Johny Wiley, New Y ork 1974, VoL 5,

Chap.9 37. Getoff, N., Schenk, G.O.: Radiation Res. 31, 486 (1967) 38. Gieskes, J.M.: in The Sea (Goldberg, E.D. Ed.). Wiley-Interscience, New York 1974, VoL5,

Chap.3 39. Goguel, J.: C. R. Aead. Sei. Paris 287, Serie D-333 (1978) 40. Goudriaan, J., Ajtay, G.L.: in The Global Carbon Cyde (Bolin, B., Degens, E.T., Kempe, S., Ket-

ner, P. Eds.) SCOPE Rep. 13. John Wiley, Chichester 1979, Chap.8 41. Hammann, R., Ottow, J.C.G.: Z. Pflanzenern. Bodenk 137, 108 (1974) 42. Higgins, I.J.: Mierobial Bioehern. 21, 300 (1978) 43. Hodgson, G.W., Ponnamperuma, C.: Proe. NatL Aead. Sei. (US), 59, 22 (1968) 44. Holland, H.D.: The Chemistry of the Atmosphere and Oceans. John Wiley, New York 1978 45. Honjo, S.: in The Fate of Fossil Fuel CO2 in the Oeeans (Andersen, N.R., Malahoff, A. Eds.).

Plenum Press, New York 1977, pp. 269-294 46. Hutehinson, G.E.: A Treatise on Limnology. J. Wiley, New York 1957, VoLl 47. Ingle, S.E.: Marine Chem. 3, 301 (1975) 48. Ishimoto, M., Egami, F.: Proe. First Internat. Symp. Origin of Life on Earth (Clark, F., Synge,

R.L.M. Eds.). Pergamon Press, New York 1957, pp. 555-561 49. Jannaseh, H.W., Wirsen, C.O., Taylor, C.D.: AppL Environ. MicrobioL 32, 360 (1976) 50. Keeling, C.D.: J. Geophys. Res. 73,4543 (1968)


51. Keeling, C.D.: in Chemistry of the Lower Atmosphere (Rasool, S.l. Ed.). Plenum Press, New York 1973, Chap. 6

52. Keeling, C.D. et al.: TeJlus 28,538 (1976) 53. Keeiing, C.D. et a!.: ibid. 28, 552 (1976) 54. Keeiing, C.D., Baeastow, P.B.: in Energy and Climate, Stud. Geophys., 72-95, US-Nat. Aead.

Sei., Washington, D.C. 1977 55. Kelts, K., Hsü, K.J.: in Lakes-Chemistry, Geology, Physics (Lerman, A. Ed.). Springer, New

Y ork 1978, Chap.9 56. Kenyon, D.H., Steinmann, G.: Biochemical Predestination. McGraw-Hill, New York 1969 57. Kester, D.R., Pytkowicz, R.M.: in Global Chemieal Cyc1es and their Alterations by Man (Stumm,

W. Ed.). Dahlem Konferenzen, Berlin 1977, pp. 99-120 58. Konopka, A., Brock, T.D.: Appl. Environ. Mierobiol. 36, 572 (1978) 59. Kroopniek, P., Craig, H.: Earth Planet Sei. Lett. 32, 375 (1976) 60. Kusnezow, S.l.: The Mieroflora of Lakes and its Geoehemieal Activity. Academy Nauk, Lenin

grad 1970 (in Russian) 61. Lemmon, R.M.: Chem. Rev. 70, 95 (1970) 62. Lerman, A.: in Global Chemical Cyc1es and their Alterations by Man (Stumm, W., Ed.). Dahlem

Konferenzen, Berlin 1977, pp. 275--289 63. Li, Y-H., Takahashi, T., Broecker, W.S.: J. Geophys. Res. 74, 5507 (1969) 64. Likens, G.E. et al.: Biogeochemistry of a Forested Ecosystem. Springer, New York 1977 65. Machta, L.: in The Changing Chemistry of the Oceans, Nobel Symp. 20 (Dryssen, D., Jagner, D.

Eds.) Almquist and Wiksell, Stockholm 1972, pp. 121-145 66. Manabe, S., Wetherald, R.T.: J. Atmos. Sei. 32, 3 (1975) 67. Martens, C.S., Berner, R.A.: Limnol. Oceanogr. 22, 10 (1977) 68. Martin, D.F.: Marine Chemistry. Marcel Dekker, New York 1970, Vol.2, Chap.8 69. Mathur, S.P., Paul, E.A.: Nature 212,646 (1966) 70. McInemey, M.J., Bryant, M.P., Pfennig, N.: Arch. Microbiol. 122, 129 (1979) 71. Megard, R.O.: 1968, cited by Kelts and Hsü [55] 72. Menzel, D.W.: in The Sea (Goldberg, E.D. Ed.) John Wiley, New York 1974, Vol. 5, Chap.18 73. Miller, S.L.: Science 117, 528 (1953) 74. Miller, S.L., Urey, H.C.: Science 130, 245 (1959) 75. Mopper, K., Degens, E.T.: in The Global Carbon Cyc1e (Bolin, B., Degens, E.T., Kempe, S., Ket

ner, P. Eds.) Scope Rep. 13. John Wiley, Chichester 1979, Chap.ll 76. MueJler, G.: in Encyc10pedia of Science and Technology. McGraw-HiJl, New York 1971, Vol.2,

p.25l 77. Nissenbaum, A., Kaplan, l.R.: Limnol. Oceanogr. 17, 570 (1972) 78. Ohle, W.: Arch. Hydrobiol. 46, 153 (1952) 79. Oparin, A.l.: Proiskhozhenie zhizni. Izd. Moskovskii Rabochii, Moscow 1924 80. Ottow, J.c.G.: Naturwissenschaften 65, 413 (1978) 81. Park, K.: Science 162, 357 (1968) 82. Parsons, T.R.: in Chemical Oceanography 2nd ed. (Riley, J.P., Skirrow, G. Eds.). Academic Press,

London 1975, Vo1.2, Chap. 13 83. Peng, T.-H. et al.: in Fate ofFossil Fuel COz in the Oceans (Andersen, N.R., Malahoff, A. Eds.).

Plenum Press, New York 1977, pp. 355-373 84. Pfennig, N.: Ann. Rev. Microbiol. 21, 285 (1967) 85. Pfennig, N., Trüper, H.G.: in Bergey's Manual ofDeterminative Bacteriology 8th ed. (Buchanan,

R.E., Gibbons, N.E. Eds.). Williams and Wilkins, Baltimore 1974, Part 1 86. Platt, T., Subba, Rao, D.V.: Fisheries Res. Board Canada, Techn. Rep. No. 370 (1973) 87. Ponnamperuma, C. et al.: Proc. Nat!. Acad. Sei (US) 49, 737 (1963) 88. Ponnamperuma, c., Woeller, F.H.: Curr. Modern Bio!. 1, 156 (1967) 89. Reeburgh, W.S.: Earth Planet. Sei. Lett. 28, 337 (1976) 90. Reichle, D.E. et al.: in Carbon and the Biosphere (Woodwell, G.M., Pecan, E.V. Eds.) 24th

Brookhaven Symp. Biology. USAEC Conf.-720510, pp. 345-365 (1973) 91. Reid, c., Orgel, L.E., Ponnamperuma, c.: Nature 216, 936 (1967) 92. Reiners, W.A.: in Carbon and the Biosphere (Woodwell, G.M., Pecan, E.V. Eds.) 24th Brook

haven Symp. Biology. USAEC Conf.-720510, pp. 345-365 (1973) 93. Reiners, W.A.: ibid. pp. 368-382

The Carbon Cycle 109

94. Richards, F.A., et a!.: Limnol Oceanogr. 10, (Supp!.), R 185 (1965) 95. Riley, G.A.: in Carbon and the Biosphere (Woodwell, G.M., Pecan, E.V. Eds.) 24th Brookhaven

Symp. Biology: USAEC Conf.-720510, pp. 204-220 (1973) 96. Ring, D. et a!.: Proc. Nat!. Acad. Sci. (US) 69,765 (1972) 97. Rosswall, T.: Nitrogen, Phosphorus, and Sulphur-Global-Cycles (Svensson, B.H., Söderlund, R.

Eds.). SCOPE Rep. 7, Eco!. Bu1l22, 157 (1976) 98. Rotty, R.M.: 1977, cited by Freyer [34] 99. Sehlesinger, W.H.: Ann. Rev. Eeo!. Syst. 8, 51 (1977)

100. Schneider, S.: J. Atmos. Sci. 32, 2060 (1975) 101: Schnitzer, M., Khan, S.U.: Humic Substances in the Environment. Marcel Dekker, New York

1972 102. SCOPE-Rep. No. 13, The Global Carbon Cycle (Bolin, B., Degens, E.T., Kempe, S., Ketner, P.

Eds.). John Wiley, Chichester 1979 103. Sekihara, K.: in Environmental Chemistry (Bockris, J. O'M. Ed.). Plenum Press, New York 1977,

Chap.1O 104. Siegenthaler, U., Oeschger, H.: Science 199, 388 (1978) 105. Skerman, V.B.D., Mac Rae, I.c.: Can. J. Microbio!. 3, 505 (1957) 106. S0rensen, J.: App!. Environ. Microbio!. 35, 301 (1978) 107. Sorokin, Y.I.: Arch. Hydrobio!. 66, 391 (1970) 108. Strong, A.E., Eadie, B.J.: Limno!. Oceanogr. 23, 877 (1978) 109. Stumm, W., Morgan, J.J.: Aquatie Chemistry. John Wiley, New York 1970 110. Stumm, W. (Ed.): Global Chemical Cycles and their Alterations by Man. Dah1em Konferenzen,

Berlin 1977 111. Sybesma, Chr.: in Photobiology of Mieroorganisms, (HalIdai, P. Ed.). Wiley-Interscience, Lon

don 1970, Chap.3 112. Takahashi, H., Tanaguchi, S., Egami, F.: in Comparative Biochemistry (Florkin, M., Mason,

H.D. Eds.) Academic Press, New York 1963, pp. 91-202 113. Thauer, R.K., Jungermann, K., Decker, K.: Bacterio!. Rev. 41, 100 (1977) 114. Trimble, R.B., Ehrlich, H.L.: Baet. Proc. 1968, 135 115. Turner, J.S., Brittain, E.G.: Bio!. Rev. 37, 130 (1962) 116. Urey, H.C.: Proc. Nat!. Acad. Sci. (US) 38, 351 (1952) 117. Valientyne, J.R.: in Encyclopedia of Science and Teehnology. McGraw-Hili, New York 1971,

Vo!. 2, p.251 118. Wetzei, R.G.: Verh. Internat. Verein Limno!. 16, 321 (1966) 119. Wetzei, R.G., Allen, H.L.: in Productivity Problems ofFreshwaters (Kajak, Z., Hillbricht-Ilkow

ska, A. Eds.) PNW Polish Sci. Publishers. Warsaw, 1970, pp. 333-347 120. Wetze!, R.G., Rieh, P.H.: in Carbon and the Biosphere (Woodwell, G.M., Peean, E.V. Eds.). 24th

Brookhaven. Biology USAEC. Conf.-72051O, pp. 241-263 (1973) 121. Wetzei, R.G., Otsuki, A.: Arch. Hydrobio!. 73, 15 (1974) 122. White, W.S.: Role of Calcium Carbonate in Lake Metabolism. Ph. D. Thesis, Michigan State

Univ. 1974 123. Whittaker, R.H., Likens, G.E.: in Carbon and the Biosphere (Woodwell, G.M., Pecan, E.V. Eds.).

24th Brookhaven Symp. Biology, USAEC Conf-720510, pp. 281-302 (1973) 124. Whittaker, R.H.: Communities and Ecosystems, 2nd ed. Macmillan, Toronto 1975

125. Whittaker, R.H., Likens, G.E.: in Primary Productivity of the Biosphere (Lieth, H., Whittaker, R.H. Eds.). Springer, New York 1975, pp. 281-302

126. Withrow, R.B.: in Photoperiodism phenomena in plants and Animals. Pub!. Am. Ass. Advrnt. Sci. 55, 439 (1959)

127. Williams, P.M. Oeschger, H., Kinney, P.: Nature 224, 256 (1969)

128. Williams, P.J. le B.: Chemical Oceanography, 2nd ed. (Riley, J.P., Skirrow, G. Eds.). Academie Press, London 1975, Vo1.2, Chap. 12

129. Woese, C.R.: J. Mol. Evol. 13,95 (1979) 130. Wolman, Y., Haverland, W.J., Miller, S.L.: Proc. Natl. Acad. Sci. (US) 69, 809 (1972) 131. Woodweli, G.M., Pecan, E.V. (Eds.): Carbon and the Biosphere, 24th Brookhaven Symp. Biology

USAEC Conf.-72051O (1973) 132. Woodwell, G.M., Rieh, P.H., Hall, C.A.S.: ibid. pp. 221-240

110 A. l. B. Zehnder

133. Woodwell, G.M., Houghton, R.A.: in Global Chemical Cyc1es and their Alterations by Man (Stumm, W. Ed.). Dahlem Konferenzen, Berlin 1977, pp. 61-72

134. Woodwell, G.M.: Sei. Am. 238, 34 (lan.) (1978) 135. Woodwell, G.M. et al.: Science 199, 141 (1978) 136. Zafiriou, O.c.: Marine Chem. 5, 497 (1977) 137. Zehnder, A.l.B.: in Water Pollution Microbiology (MitchelI, R. Ed.). John Wiley, New York

1978, Vo1.2, Chap.13 138. Zehnder, A.J.B., Brock, T.D.: J. Bacteriol. 137,420 (1979) 139. Zelitch, 1.: Photosynthesis, Photorespiration and Plant Productivity. Academic Press, New York

1971

Molecular Organie Geochemistry

P. A. Schenck, J. W. de Leeuw

Delft University of Technology Department of Chemistry and Chemical Engineering Organic Geochemistry Unit Delft, The Netherlands

Introduction

Organic geochemistry can be characterized as a field of science which studies the "fate" of organic compounds in sediments. In its earlier days research efforts were limited - from sheer lack of other possibilities - to the determination ofbulk characteristics as for instance contents of organic carbon, nitrogen, hydrogen and percentage of organic matter extractable with organic solvents.

Development of chromatographie methods in general and of gas chromatography and gas chromatography/mass spectrometry more specifically have stimulated organic geochemical work enormously over the past two decades. As a result studies have become possible on the molecular level thus leading to a much better insight into the processes occurring in sediments and involving organic compounds. Consequently much more has become known about the relation between organic compounds found in recent or ancient sediments and crude oils on the one hand and their precursors in living nature on the other hand.

This knowledge may be of help in environmental studies since it gradually becomes better known which specific compounds and what amounts can be expected as natural background and/or input contrary to what has to be considered as result ofhuman activities [1]. The present chapter does not attempt to cover the complete fie1d of organic geochemistry. On a limited number of c1asses of compounds it will be demonstrated what detailed insight can be obtained in several cases with our present day analytical techniques. Because of the importance of hydrocarbons from crude oils as pollutants in the environment, several groups of hydrocarbons and their possible natural precursors have been chosen to illustrate the development of organic geochemical knowledge.

112 P.A. Schenck and J. W. de Leeuw

Normal Alkanes

The occurrence of n-alkanes in organisms, sediments and crude oils has been investigated extensively [1, 2]. These studies are facilitated by the fact that gas chromatographie and mass spectrometric methods make their quantitation relatively simple [3].

It has been known for a long time that crude oils contain very differing amounts of n-alkanes, varying form dominant in certain fractions to undetectable amounts [2, 4]. In paraffinic crudes significant differences in distribution may occur, although in the molecular weight range over n-C25 the relative amounts are always decreasing with increasing chain length.

The study of n-alkanes in extracts from recent and ancient sediments has attracted much interest since the finding of recent hydrocarbons in recent sediments by P. V. Smith [5] and even more after Stevens, Bray and Evans [6] showed that the distribution ofn-alkanes in recent sediments is significantly different from that in crude oils. The consequence of this latter finding is that oil accumulations are not the result of mere concentration of hydrocarbons already present. In other words, there must be an oil genereation process. A relatively great number of recent sediments investigated are characterized by a preferent occurrence ofthe odd numbered n-alkanes with chain lengths greater than 15 carbon atoms. Tbis odd predominance is most characteristic in the range n-C25 - n-C33 . The odd predominance is often expressed by means of the carbon preference index (C.P.I.) defined as [7]:

C P I =!{C25 +C27+C29+C31 +C33 C25 +C27+C29+C31 +C33 } ... 2 C24+C26+C28+C30+C32 + C26+C28+C30+C32+C34

In many cases use is made of the R 29 value [8J, defined as R 29 = C 2 X C~9 , 28+ 30

because the n-C29 is often the most predominating n-alkane and because the calculation in the latter case is much simpler.

Numerous studies have revealed that an odd predominance in tbis range is caused by the presence ofn-alkanes belonging to higher terrestrial plants [9,10] in which they are biosynthesized [11]. Most of the recent sediments investigated have an input of terrestrial material brought into them by rivers. Marine organisms on the other hand are known for their preference for one specific n-alkane, mostly n-C15, n-C17 or n-C19 [12]. Consequently, recent sediments containing only marine organic matter do not show the above mentioned odd predominance in the high molecular weight range. The odd predominance gradually disappears [2] in sediments that have undergone the influence ofhigher temperatures because n-alkanes without any preference are formed from organie material in the sediments. Tbis process is known as maturation. The n-alkanes originally present are diluted by those formed during diagenesis ofthe organic matter. This brings the C.P.I. value in crude oils down to about l.0-l.3. In some cases, when sediments have not been exposed to higher temperatures in the earth crust, the original n-alkane distribution pattern can still be recognised. A relevant example is given by Knoche and Ourisson [13] for the n-alkanes isolated from arecent and a fossil (50 x 106 yrs old)

Molecular Organic Geochemistry 113

plant (Equisetum) in which the odd numbered n-alkanes C23 , C25 , C27 , C29 and C31 dominate in a comparable way.

The characteristic distribution of n-alkanes in recent sediments offers possibilities for discriminating between fossil and recent alkanes. It should be kept in mi nd that the fossil alkanes can be either anthropogenie or brought into the sediment by natural influences (seeps, erosion). The choise between these two explanations has to be made on other grounds than the geochemical considerations mentioned here, like e.g. the actual presence of natural oil seeps in the area under consideration. A disturbing factor when using n-alkane distributions for determining anthropogenie origin and even a specific origin of pollution (e.g. emde oil spills) is the fact that n-alkanes are relatively easily consumed by microorganisms, causing changes in composition after the emde oil has entered the environment [14]. Although n-alkane distribution patterns may be of great he1p in finding sources of pollution, other data and considerations (e.g. period of time in the environment, analytical control of possible sources) have to be taken into account.

Acyclic Isoprenoid Hydrocarbons

The observation ofthe acyclic isoprenoid hydrocarbons pristane (2,6,1O,14-tetramethyl pentadecane) and phytane (2,6,10, 14-tetramethyl hexadecane) in sediments and emde oils [15, 16] has focussed interest on the group of acyclic polyisoprenoids. They can be considered as significant examples of biochemical fossils because of their stmctural relationship to naturally occurring compounds containing the same carbon skeleton. A direct relation between phytane and phytol has been suggested early [17], phytane thus being an indicator for the original presence of photosynthetic organisms. Phytol is li be ra ted from chlorophyll in a very early stage of degradation; pristane and phytane might weIl be formed from it via aseries of complex conversions [18, 19]. Stimulation experiments in the laboratory have given proof for several reactions mentioned in the scheme by Didyk et al. [20].

The interrelation between phytol and the isoprenoid hydrocarbons Cl9 and C20 has been shown by studying the stereochemistry of the corresponding chiral centra in phytol and in the isoprenoid hydrocarbons respective1y [21-24]. It could proved to be the same in recent sediments [24]. It has been shown, however, from both simulation experiments [25] and from results of analyses of extracts from sediments of increasing "maturity" [26] that epimerization at the chiral centra in the isoprenoid hydrocarbons occurs with increasing degree of diagenesis.

Apart from the above-mentioned isoprenoids pristane and phytane aseries of isoprenoid hydrocarbons of lower molecular weight have also been found [27]. They are often considered as being formed from phytol by exclusive cleavage of carbon bonds of the main chain. This mere cleavage of the main chain's carbon bonds could provide an explanation for the fact that the Cl7-isoprenoid (2,6,10-trimethyl tetradecane) as well as the C12 (2,6-dimethyl decane) are often virtually absent in extracts from sediments or emdes or only present in minor amounts [28].

On the other hand, lower molecular weight isoprenoid hydrocarbons like farnesane (2,6, lO-trimethyl dodecane) could also be derived from corresponding alcohols like farnesol. I t is interesting in this respect that in two emde oils from the Gulf


Coast and Syria all four famesane stereoisomers occur in equivalent amounts [29]. This may point indeed to farnesol as aprecursor although phytol as precursor cannot be precluded in case of epimerization at the chiral centra.

Another view on the origin of the isoprenoid hydrocarbons is presented by their occurrence as such in recent organisms. Blumer and his collaborators showed as early as the mid-sixties that phytol can be converted into pristane by copepods (or bacteria living in their intestinal system [30, 31]. Consequently pristane can thus be introduced into recent sediments as such. Unsaturated hydrocarbons with the same carbon skeleton could be isolated too [32]. It is noteworthy that no phytane could be found. Recent work has made it clear that many more isoprenoid hydrocarbons occur as such in the so called Archaebacteria; the whole range of CIS-C30

could be shown in several species, most of these isoprenoids of the regular headto-tail type [33]. Since part of the methane producing bacteria belong to these Archaebacteria it may very weIl be that at least part of the isoprenoid hydrocarbons found in recent or ancient sediments and in crude oils have been introduced as such and have not been derived via diagenetic pathways from phytol or comparable isoprenoid alcohols. In this context it should be emphasized that the Cl7-isoprenoid hydrocarbon has been observed in some Archaebacteria, viz. in several thermoacidophilic species [33]. Considering what has been said above on the cleavage ofthe main carbon chain of phytol it could be speculated that this Cl 7-isoprenoid hydrocarbon enters the sediment as such and is not the product of diagenetic reactions of phytol.

In addition to the free isoprenoids several species of Archaebacteria contain among others dialkylether glycerides with phytanyl groups as alkyl moieties [34, 35]. These ethers might undergo diagenetic reactions leading to the formation of isoprenoid hydrocarbons. The stereoehemistry of these phytanyl groups is the same as that of phytol as far as the two corresponding chiral centra are concemed [36]; determination of the stereochemistry at these centra thus does not point unambiguously to an origin from phytol.

The isoprenoids up to C20 mentioned hitherto all belong to the so called "regular", i.e. "head-to-tail" type. Many "irregular" ones of both the "head-to-head" and "tail-to-tail" type are found with more than 20 carbon atoms. An example of this group is squalane (2,6,10,15, 19,23-hexamethyl tetracosane) found in sediments and erude oils [37]. It eontains a "tail-to-tail" eombination; it is almost eertainly related to squalene, widely occurring in living nature.

Isoprenoid struetures with "head-to-head" bound isoprene units are also known in living nature: De Rosa et al. [38] found di-(biphytanyl) diglycerol tetraether lipids in some thermoacidophilic Archaebacteria with w-wl-diphytanyl structures. This finding is the more interesting sinee this same type of biphytanyl struetures is present in some kerogens, the insoluble part of organie matter in sediments [39,40].

In eonclusion it ean be stated that at the present stage of our knowledge isoprenoid hydrocarbons in (ancient) sediments and crude oils are not necessarily mainly derived from phytol from chlorophyll. On the eontrary, it might very weIl be that a eontribution from Archaebacteria plays a mueh more important role than previously antieipated; the hydroearbons ean be introdueed into the sediment either as such and/or derived from the bacterial membranes (Fig.l).


Fig.1

Chlorophyll

! Phytol

C-O~ t-o~ I C-OR

~/ Isoprenoid hydrocarbons

l~p~~1d / ~ Hydrocarbons y-O~-Y in Archaebacteria C-O~-y

I C -OR RO -C

Steroids

Occurrence and Diagenesis

115

Steroids are tetracyclic isoprenoid compounds which occur widespread in nature. Almost all eukaryotic organisms contain free and bound steroids as lipid components. The steroids present in eukaryotic organisms have specific distribution patterns as the result ofbiosynthesis and/or dietary uptake. Only a few prokaryotic organisms contain steroids. The steroids biosynthesized by these prokaryotic organisms are often of a different structural type (e.g. 4,4-dimethyl substitution) when compared with those biosynthesized by eukaryotic organisms [41]. Numerous detailed investigations of steroid components oceurring in different types of sediments ranging from very recent to very aneient have enabled organie geoehemists to unravel the major ehemieal pathways operating in sediments shortly after deposition of fresh organic matter and during maturation of sediments.

Figure 2 shows a simplified overview of the fate of naturally occurring sterols after burial in sediments. The sterols present in organisms are free or bound (e.g. as steryl ester, sulfate ester, ete.). The bound sterols have to be hydrolyzed to the free ones before they ean undergo bioehemieal transformation reaetions in the top layers of the sediment [42]. Free sterols (I) are transformed by two different pathways. Direet dehydration results in the formation of,d 3,5 -steradienes (11) and other isomerie steradienes, whenever the substrate possesses the eommon ,d5-double bond [43]. Naturally occurring 5ocH-stanols are probably dehydrated to ,d2-5ocHsterenes (111). Another bioehemieal pathway involves the transformation of ,d5_ sterols to either 5ocH- or 5ßH-stanols via several intermediates such as ,d4-stenones and saturated stanones [44, 45]. 5ßH-Stanols are only generated in this way by anaerobic microorganism [42]. The stanols are eonverted to the corresponding ,d2_ sterenes (III) and to the ,d4_ and ,d5-sterenes (IV). The ,d4_ and ,d5-sterenes mayaiso be the result of isomerization of initially formed ,d2-sterenes [46]. All these transformation reaetions take plaee in the very early stages of diagenesis and are mainly of microbiological nature as incubation studies have shown. Starting from these early stage diagenesis produets (11,111, IV) several pathways during further matu-

c Q

HO

R

A

n

~i ~

XI

.j

H

R

25

I \

via

tf -s

teno

nes

ond

50 e

r 5ß

H-s

tono

ls

501 ~

-H20

R

R

~

III

5Ybockb~

/ re

orr

.\

R

R

A

Y (

50. o

r5ßH

,17a

.H,2

0R)

3ZI

(20R

+20

S)

eo

rly

stag

e di

agen

esIs

(mo

lnly

b

,och

em

lca

l)

I R

~~

~

w

R·'A

~~

XII

R

:lZIII

R=

H,M

e,E

t

R' =

H,

Me

R-=

H,

Me

H

(50.

H,lI

.ßH

, 17ß

H; 2

0R+2

0S)

Fig

.2

(13a

.H,1

7ßH

and

13ßH

,17a

.H;2

0R+S

)

R

-0\ :-c ?> CI> I [ :-- ~ ~ i


ration can occur. Part ofthe sterenes and steradienes oftype II, III, and IV are hydrogenated to 5a- or 5ß-steranes (V) with preservation of the natural stereochemistry at positions 14, 17, and 20 (14aH, 17aH, 20R). Upon increasing maturation isomerization reactions occur resulting in more thermostable steranes, such as SaH - and 5ßH -steranes with the 14ßH - and 17 ßH -configuration and both the 20R and 20S stereoisomers (X) [47]. Another part of the steradienes of type II may undergo A-ring aromatization with either loss or a shift of the original angular methyl group at C 1D (XI). Upon further maturation the aromatization can spread out from the A-ring via the B-ring to the C-ring, ultimately resulting in a suite of completeley aromatized steroids of type XII.

Apart from saturation the ,14_ and ,1 5-styrenes (IV) undergo backbone rearrangements resulting in the so called diasterenes of type VI with 20R and 20S stereoisomers [46]. These diasterenes are either hydrogenated to isomerie mixture of diasteranes (VIII) or they are aromatized starting from ring C (VII) to completely aromatized steroids of type IX with loss or shift of the original angular methylgroups at C 1D and C13 [48, 49].

In summary we can say that in ancient sediments or oils the original sterols are reflected by isomerie mixtures of saturated steranes of types VIII and X and aromatized steroids of types XII and IX. The relative distribution of these four groups of steroid hydrocarbons is a result of all kind of sediment parameters, such as anoxie vs oxic deposition, the mineral and eIay content, the water content and the thermal gradient.

Especially the nature ofthe inorganic matrix probably plays an important role in the ultimate ratio of regular steranes and back bone rearranged steranes. It has been shown that so called superacid sites, present in eIay minerals such as kaolinite and montmorillonite, catalyze the backbone rearrangement transformations [50].

Steroids as Biological Markers

As shown above, sterols undergo a variety of diagenetic reactions. However, detailed structural elucidation of the sterol derived components present in sediments still gives eIues to the original environment of deposition. Especially the structure and stereochemistry of the several side chains and their distribution pattern enable us to reconstruct to some extent the original environment. In general it can be stated that marine organisms, especially algae, contain complex mixtures of sterols with a great variety in side chain structures. Higher plants on the contrary exhibit simple sterol patterns and the sterol side chain structures are generally limited [44]. In marine algae the number of carbon atoms of the steroidal side chain varies from 2 to 11, in higher plants form 8 to 10 only. Within the overlapping Cs to C 1D side chain series discrimination of an origin from either algae or higher plants is possible to some extent due to specific structures and the stereochemistry. Dinoflagellates and to some extent coelenterates and diatoms for instance biosynthesize sterols with 23,24-diMe substitution [51-53], while higher plants exeIusively biosynthesize sterols with side chains having alkyl substituants (Me or Et) at C24.

Moreover the stereochemistry of the alkyl substituents at C24 differs between higher plants and algae [41]. Among steroids obtained from sediments aseries of


4-methyl substituted steroid hydrocarbons is often encountered. As far as it is known only dinoflagellates and a limited number of methanotrophic bacteria are capable of 4-methylsterol biosynthesis. The occurrence of 4-methyl steranes in combination with typical algal side chain structures such as 23,24-diMe substitution in sediments reflects therefore the original presence of dinoflagellates [54]. By means of advanced capillary gas chromatography it is possible to separate 24R and 24S isomers [55]. Application ofthis method to mixtures of steranes obtained from sediments will enhance the value of steroid hydrocarbons as molecular fossils.

When crude oils or ancient sediments become available for microbial attack several phenomena have been observed as far as the steranes are concerned. Moderate biodegradation hardly affects the steranes, but heavy biodegradation results in destruction of regular steranes and survival of the diasteranes (20R better than 20S) [56].

Due to the rather detailed knowledge of steroid diagenesis and due to the taxonomic specificity of the sterol biosynthesis, it can be stated in conclusion, that a detailed study of sedimentary steroid hydrocarbons enables us to estimate the degree of maturation of a sediment and to partly reconstruct the paleo-environment.

Triterpenoids

Occurrence and Diagenesis

Introduction

The triterpenoids are widespread both in the biosphere and the geosphere [57, 58]. A relatively large number of organic geochemical investigations (especially those by the Organic Geochemistry Unit ofthe Louis Pasteur University at Strasbourg, France) have resulted in a detailed knowledge of the pathways by which several classes ofnaturally occurring triterpenoids are converted during early and late diagenesis. Since these diagenetic pathways - which determine the fate oftriterpenoids - are highly dependent on structural entities present or absent in the starting compounds it is nescessary to chemically classify the naturally occurring triterpenoids in several groups. Figure 3 schematically shows the several triterpenoid classes which will be described separately further on in this paragraph.

For geochemical reasons the triterpenoids can be divided firstly into triterpenoids having the hopanoid skeleton and those with other skeletons. The non-hopanoid triterpenoids occur widespread in higher plants and very often possess an oxygen function at the C3 position (either alcohol or ketone). These non-hopanoids do not have representatives with more than 30 carbon atoms, in other words no extended homologues occcur. The hopanoids can be divided into two subgroups: the 3-oxy hopanoids and the 3-desoxy hopanoids.

The 3-oxy hopanoids occur in a few families ofhigher plants. Finally the 3-desoxi hopanoids can be divided into the so called non-extended and extended ones. The non-extended 3-desoxy hopanoids do have carbon skeletons with a


Triterpenoids

A hoponoids non-hoponoids (3-ox Yinon-extended)

(moinly higher plonts)

e.g.; s~eleton: 21

i .,(

i

~

A ~

•..

• i

HO

3-oxy-hoponoids 3-desoxy-hoponOids (few fomilles of higher plonts) e.g. :

/. ,~[ O~ extended

119

non-extended

(ferns,lichens,some microorgonisms) e.g.;

(several groups of microorganismsl

~ Fig.3

maximum of 30 carbon atoms. They mainly occur in ferns, lichens and some microorganisms. It should be noted that especially ferns do also contain other series of triterpenoids, such as fernene, adianene and filicene derivatives.

The extended 3-desoxy hopanoids exclusively occur in certain groups of microorganisms and are characterized by an extended side chain of 8 carbon atoms at position 21 often substituted with 4 or 5 hydroxygroups [58]. After this classification of the naturally occurring triterpenoids in terms of structural differences and related occurrences in organisms an overview is given of the several diagenetic pathways and the subsequent resulting suite of geochemical products for each defined class of triterpenoids.

3-0xy-Triterpenoids (N on-Hopanoids)

The triterpenoids belonging to this dass which occur alm ost exdusively in higher plants are converted in the first stages of diagenesis. Two major diagenetical pathways can be discriminated (Fig.4; ß-amyrin as an example).

120 P. A. Schenck and J. W. de Leeuw

Dicgenesis of 3-cxy-triterpenoids (non-hopcnoids)

?? ) )crom.

Lass of A_ring (via ~l crom. HOf

:n:

Fig.4

It is thought that microorganisms play an important role in these transformation reactions, although it cannot be precluded that the initialloss of the original A-ring (pathway 11) occurs via a photochemical conversion [59]. The first products via pathway I are the AB or ABC-ring aromatized compounds depending on the starting product. Possible intermediates with one aromatized ring have not been found. Obviously, during the aromatization several methylgroups are removed [60, 61]. During further diagenesis the aromatization continues via intermediates with ABC-aromatic rings to the fully aromatized compounds. Similar end products are known with more or less methylgroups [61]. Following pathway 11 the A-ring is lost, possibly via the intermediate shown in Fig.4 which points to microbial activity. The resulting products possess an aromatic B-ring. Ultimately, due to further aromatization of the C-, D-, and E-ring, fully aromatized tetracyclic end products are generated [61, 62]. Due to the original precursor molecule and due to slightly different diagenetic pathways the methyl substitution pattern varies to some ex-


Diogenesis of 3-desoxy hoponoids

Fig.5

teht. It should be emphasized that these partly and/or fully aromatized compounds are already present in very recent sediments, indicating that these diagenetic pathways occur under very mild conditions and are probably in part of microbial nature.

It seems surprising that fully saturated triterpenoid compounds of this type are less frequently encountered in both recent and ancient sediments [59], since other cyclic compounds possessing a hydroxylgroup in the A-ring such as sterols, are partly reflected in sediments as saturated hydrocarbons, e.g. steranes.

3-Desoxy Hopanoids (Non-Extended)

In ferns, lichens and several micro-organisms non-extended 3-desoxy hopanoids occur sometimes together with other classes of triterpenoids. Figure 5 shows the several diagenetical pathways by which the 3-desoxy hopanoids, such as diplopterol are converted. During early diagenesis important intermediates are generated, e.g. the C27-ketone and C30-olefins [63]. It is believed that the intermediate ketones initiate the aromatization of ring D upon further diagenesis [60]. Starting

122 P.A. Schenck and J.W. de Leeuw

Dlageneslc; of C3S-hopanepolyols

#~H t early stage dia genesis

t

+

L arom.

~~7ßHI' 1

rr. ~ \ : '(17C1H)

+ n=O-6

Fig.6

from the D-ring aromatized eompound, rings C and Bare aromatized respeetively. The fuHy aromatized eompounds are produeed upon further diagenesis [62, 64].

Another part ofthe C2rketone-intermediate is eonverted via several intermediates to the eorresponding hydroearbon with the thermostable 17IXH -eonfiguration [63]. The initiaHy formed C30-o1efins undergo hydrogenation resulting in isomerie C30-hopanes with both the 17ßH, 21IXH and 171XH, 21ßH eonfigurations [63].

Extended Hopanoids

The C35-hopanepolyols, whieh are exdusively eneountered in a vast number of baeteria and eyanobaeteria undergo rapid ehanges during very early diagenesis (Fig. 6). A suite of intermediate eompounds with C27 , C29-C35 earbon atoms and


3-oxy-hopanoids

Fig.7 ~ ~ ···· .. fOH ,

~ o ..... , hydroxyhopanone

several functional groups have been isolated from very recent sediments [65]. Upon further diagenesis these compounds are converted via two major diagenetical pathways which parallel the pathways described for the non-extended 3-desoxy-hopanoids. Hence, the intermediates partly react to compounds with an aromatic Dring and subsequently to the ABCD-ring aromatized products via compounds with CD- and BCD aromatized rings respectively. On the other hand the initial diagenetic compounds are reduced to the corresponding hydrocarbons with the 17 ßH, 21ßH configuration. These hydrocarbons isomerize in a later stage of diagenesis to hydrocarbons with the 17aH-configuration (the C27 compound) and to hydrocarbons with the 17aH, 21ßH and 17ßH, 2laH configurations (the C29-C35 series) [63]. The latter two series of compounds occur as 22R and 22S isomers after considerable diagenesis. The recent finding of 17aH-hopanoids in a contemporary unpolluted mud suggests that 17aH-hopanoids can also be obtained by microbial processes and not only via an acid catalyzed isomerization in older sediments [65]. In a Balthic sea sediment a number of 3-methyl hopanoid hydrocarbons are shown to be present. They are of bacterial origin also, since it is known that Acetobacer xylinum and Acetobacter rancens biosynthesize hopanepolyoles with a 3-methylgroup [66].

3-0xy Hopanoids

These triterpenoids (Fig.7) occur in the resins or saponins of a few higher plant families. As yet hardly anything is known ab out the geological fate of these compounds, although one might speculate that major diagenetical pathways are comparable with those described above for the 3-desoxy hopanoids and the non-hopanoid 3-oxy triterpenoids.

Triterpenoids as Biological and Maturation Markers

A detailed analysis of triterpenoid compounds in a sediment may reveal to some extent the environment of deposition. a) The presence of extended saturated andjor aromatized hopanoid hydrocarbons

reveals the original presence of bacteria andjor cyanobacteria. b) The presence of 3-methyl-hopanoid hydrocarbons points to the original occur

rence of sepcific bacteria, like Acetobacter xylinum. c) The presence of pentacyclic and tetracyclic aromatic compounds with well de

fined structures points to the original presence of higher plant waxes and hence to a terrestrial input.


Diagenesis of di terpenoids

Fig.8

The state of maturation of the sediment can be estimated by detailed analyses of the sedimentary triterpenoids. a) Functionalized triterpenoids only occur in very recent sediments. b) The configurations of the Cl?' C21 and C22 positions in the hopanoid hydro

carbons reflect the state of relative maturity of the sediments. It should be no ted that the occurrence ofpartly or fully aromatized triterpenoid

hydrocarbons is not restricted to ancient sediments. Certain types already occur in very recent sediments.

Diterpenoids

Over the past few years a number of partly and fully aromatized tricyc1ic compounds (Fig.8) with specific alkylsubstitution patterns have been encountered in young sediments [61, 67, 68]. The cooccurrence ofthese hydrocarbons with similar functionalized tricyclic compounds provides evidence for the diagenetic pathways of diterpenoids, especially for those with the abietin skeleton. Abietic acid, a compound relatively abundant in higher plant resins, especially from conifers, is thought to be the most important precursor of these sedimentary hydrocarbons. Via a number offunctionalized intermediates [67] ring C is aromatized and the carboxylgroup is removed (pathway B).


Abietatriene may result from the reduetion ofthe earbonyl group of abietie acid (pathway A). But it also oeeurs in nature [69]. Further aromatization ultimately results in the formation of the fully aromatized eompound retene.

These aromatie hydroearbons, when found in sediments are exeellent biologieal markers of resinous higher plants. Sinee these diterpenoid derived aromatie eompounds are already present in very young sediments and soils we must assurne that the transformations are mainly mierobial in nature [70].

Polycyclic Aromatic Hydrocarbons

Within the class of aromatie hydrocarbons, much attention has been paid in recent years to the polycyclic aromatic hydrocarbons. Many polycyclic aromatic hydrocarbons (PAH) are known earcinogens (e.g. benzo(a)pyrene, benz(a)anthracene, methylchrysenes). The oeeurrence of polycyclic aromatic hydrocarbons in geological sampies in the environment has therefore increased the search for these compounds in different environments.

Examples of relatively simple polycyclic aromatic hydrocarbon mixtures in geological settings are benzopyrene in asbestos [71-73]; fluoranthene and picene in mercury ores [74], perylene in recent sediments [75-78] and pyrene and fluorenthene in mangane se nodules [79].

It is known that in ancient sediments and erude oils very complex mixtures of polyeyclie aromatic compounds occur [2]. Extensive search for this type of compounds has shown that they occur widespread in soils and recent aquatic sediments. Blumer and coworkers analysed many recent marine sediments [80-82]; the same type of sediments were also analysed by other investigators [67, 83, 84]. Giger and Schaffner [85, 86] and Wakeharn et al. [61, 87, 88] concentrated on lacustrine sediments.

In erude oils the alkylhomologues within aseries are more abundant than the non-substituted parent aromatic moleeule [81, 82]. In recent sediments it is the parent moleeule whieh is the predominant one; this may offer an opportunity to decide upon anthropogenie pollution or natural oecurrenee.

E.g. Blumer [81] mentions a case in whieh a specific PAH-series (the Cn H2n - 18

series) was shown to be present in an oil spill at West Falmouth (Massachusetts); its characteristics eould be found in the marsh sediment, even five years after the spill. Other series (Cn H2n - 22 and Cn H2n - 24) did not occur in the oil, but were shown in the sediment.

The remarkable similarity in composition of the PAH's over a wide range of depositional environments, the predominance of the unsubstituted "parent molecule" of aseries and the extended alkylhomology for PAH shown in recent marine sediments lead Y oungblood and Blumer [82] to the conclusion that they cannot have been formed biochemically.

Hase and Hites [89] conclude from aseries of experiments that bacteria do not produee polycyclic aromatie hydrocarbons but rather bioaccumulate them.

Thermal PAH formation can oceur over a wide range of temperatures: high temperatures favour the unsubstituted aromatics whereas lower temperatures preserve a greater degree of alkylation. These considerations explain the highly alkylated PAH assemblages in erude oils and point to expect even more highly alkylated


mixtures for this type of compounds when formed by abiotic transformations during early diagenesis of organie matter. The composition of the mixtures found suggest a pyrolytic origin at intermediate temperatures, which Y oungblood and Blumer [82] think to be forest and prairie fires, the products ofwhich are widely spread by aeolian transport.

Wakeharn et al. [61,88] have concentrated on the search for PAH in lake sediments in rather highly industrialized and very populated areas [Lake Washington (USA) and three lakes in Switzerland]. Surface sediment layers in these lakes are strongly enriched in PAH - up to 40 times - compared to deeper layers, deposited in pre-industrial revolution sediments. This led the authors to coneIude that most of the PAH are of anthropogenie origin. Based upon detailed comparison of data obtained for sediment sampies, street dust, weathered asphalt, tire particles and automobile exhaust they coneIude that urban runoff containing street dust partieIes is possibly the major present-day source for the PAH in the lakes investigated. The authors' data suggest that asphalt particles in the street dusts may be an extremely important contribution to the PAH content of the lake sediments. On the contrary a limited number of PAH are apparently of natural, early diagenetic origin [61]. This group of compounds consists of perylene, an extended series of phenanthrene homologues, retene and pimanthrene derived from diterpenes (abietic acid and primarie acid respectively), aseries oftetra- and pentacyeIic PAH derived from pentacyeIic 3-oxy triterpenoids (e.g. of the amyrin type) and pentacyeIic PAH derived from pentacyclic 3-desoxy triterpenoids, mostly of the hopane type.

The presence of perylene has been mentioned by many authors as a predominant PAH in recent freshwater [87, 90] and in marine sediments [67, 75, 77, 91]. Perylene is present though not abundant, in soils and unaltered river sediments [67, 87] and in plankton [85, 87]. Terrestrially derived biogenic perylenequinone pigments are known in nature and are therefore suggested as possible percursors. The recent findings of perylene in Namibian Shelf sediments [78] contradicts a terrestrial origin since the input of terrestrial material into these sediments is thought to be negligible.

Wakeharn et al. [61] coneIude that early diagenesis can be an important process responsible for the formation of the special PAH mentioned before in recent lake sediments. The authors state that - based upon the abundance of this suite of natural P AH in very young lake sediments - the transformation reactions of natural compounds into these PAH must be rapid and probably mediated by micro-organisms. The last hypothesis remains untested, however, up till now.

The possible formation ofP AH from natural precursors as indicated above and mentioned in the previous sections on steroids, triterpenoids and diterpenoids has to be taken into account when interpretating data on P AH regarding their origin. Only detailed analyses and careful comparison of results will open possibilities for deciding on an origin either from natural sources or due to human activities.

Epilogue

The foregoing paragraphs have illustrated the growing insight into the relationship between organic compounds occurring in nature and those found in sediments and


erude oils. Speeific distributions and elueidation of detailed struetural features as e.g. stereoehemistry open possibilities for diseriminating between reeent and fossil eompounds.

Inereasing information on the chemotaxonomy of many organisms has given a mueh better basis for the interpretation of data obtained from analyses of sediments.

With inereasingly better methods for detailed analyses more results ean be foreseen for the years to eome. This will not only give better insights into relationships between fossil and reeent eomponents, but also open horizons for diseriminating between natural and anthropogenie origin.

References

I. BrasselI, S.c., et al.: Natural Background of Alkanes in the Aquatic Environment In: Aquatic Pollutants - Transformation and Biological Effects [Hutzinger, 0., van Lelyveld, L.H., Zoetman, B.C.J. (eds.)], Oxford, Pergamon Press 1978, pp. 69-86

2. Tissot, B.P., Weite, D.H.: Petroleum Formation and Occurrence, Berlin, Heidelberg, New York, Springer 1978

3. Eglinton, G., Murphy, M.T.J.: Organic Geochemistry, Methods and Results, Berlin, Heidelberg, New York, Springer 1969

4. Hunt, J.M.: Petroleum Geochemistry and Geology, San Francisco, W.H. Freeman 1979 5. Smith, P.V.: Science 116, 437 (1952) 6. Stevens, N.P., Bray, E.E., Evans, E.D.: Bull. Amer. Ass. Petr. Geol. 40, 975 (1956) 7. Bray, E.E., Evans, E.D.: Geochim. Cosmochim. Acta 22, 2 (1961) 8. Philippi, G.T.: Geochim. Cosmochim. Acta 29, 1021 (1965) 9. Eglinton, G., et al.: Nature 193, 739 (1962)

10. Eglinton, G., Hamilton, R.J.: The distribution of alkanes. In: Chemical Plant Taxonomy [Swain, T. (ed.)], New York, San Francisco, London, Academic Press 1963, pp. 187-217

11. Kolattukudy, P.E., Croteau, R., Buckner, J.S.: Biochemistry of hydrocarbons and oxygenated derivatives. In: Chemistry and Biochemistry ofNatural Waxes [Kolattukudy, P.E. (ed.)], Amsterdam, Oxford, New York, Elsevier 1976, pp. 294-312

12. Weete, J.D.: Aigal and Fungal Waxes. In: Kolattukudy, P.E. (ed.). toc. eit. pp. 350-418 and references cited therein

13. Knoche, H., Ourisson, G.: Angew. Chem. 6, 1085 (1967) 14. Bird, C.W., Molton, P.M.: The production of fatty acids from hydrocarbons by rnicroorganisms.

In: Topics in Lipid Chemistry. Vol. 3 [Gunstone, F.D. (ed.)], London, Elek. Science 1972 15. Dean, R.A., Whitehead, E.V.: Tetrahedron Lett. 21, 768 (1961) 16. Bendoraitis, J.G., Hepner, L.S.: Anal. Chem. 34, 49 (1962) 17. Calvin, M.: Chemical Evolution. London, Oxford University Press 1969, pp. 70-88 18. Brooks, P.W., Maxwell, J.R.: Early stage fate of phytol in a recently - deposited lacustrine sedi

ment. In: Advances in Organic Geochemistry - 1973. [Tissot, B., Bienner, F. (eds.)], Paris, Ed. Technip. 1974, pp. 977-992

19. de Leeuw, J.W., et al.: Phytol derived compounds in the geosphere. In: Advances in Organic geo-chemistry - 1975 [Campos, R., Goni, J. (eds.)], Madrid, Enadimsa 1977, pp. 61-80

20. Didyk, B.M., et al.: Nature 272, 216 (1978) 21. Cox, R.E., et al.: Chem. Comm. 1639 (1970) 22. Maxwell, J.R., et al.: The diagenesis and maturation of phytol. The stereochemistry of 2,6,10,14-

tetramethyl pentadecane from an ancient sediment. In: Advances in Organic Geochemistry 1971 [von Gaertner, H.R., Wehner, H. (eds.)], Pergamon Press 1972, pp. 277-291

23. Brooks, P.W., Maxwell, J.R., Patience, R.L.: Geochim. Cosmochim. Acta 42, 1175 (1978) 24. Patience, R.L., Rowland, S.J., Maxwell, J.R.: Geochim. Cosmochim. Acta 42, 1871 (1978) 25. Borgohain, M.: Alteration of lipids and kerogens, Thesis, Bristol 1971 26. MacKenzie, A.S., et al.: Geochim. Cosmochim. Acta 44, 1709 (1980)


27. McCarthy, E.D.: A treatise in Organic Geochemistry, Thesis, University of California 1967 28. McCarthy, E.D., Calvin, M.: Tetrahedron 23, 2609 (1967) 29. Brooks, P.W., et al.: Stereochemical studies of acyc1ic isoprenoid compounds. VI: The stereo

chemistry offarnesane from crude oil. In: Advances in Organic Geochemistry-1975 [Campos, R., Goni, J. (eds.)], Madrid, Enadimsa 1977, pp. 81-98

30. Blumer, M., Mullin, M.M., Thomas, D.W.: Science 140, 974 (1963) 31. Avigan, J., Blumer, M.: J. Lip. Res. 9, 350 (1978) 32. Blumer, M., Thomas, D.W.: Science 148, 370 (1965) 33. Holzer, G., Oro, J., Tornabene, T.G.: J. Chrom. 186,795 (1979) 34. Tornabene, T.G., et al.: J. Mol. Evol. 11,259 (1978) 35. Kates, M.: Ether -linked lipids in extremely halophilic bacteria. In: Ether Lipids [Snijder, F. (ed.)],

New York, London, Academic Press 1972, pp. 351-398 36. Anderson R., et al.: Geochim. Cosmochim. Acta 41, 1381 (1977) 37. Gardner, P.M., Whitehead, E.V.: Geochim. Cosmochim. Acta 36,259 (1972) 38. De Rosa, M., et al.: J.C.S. Chem. Comm. 514 (1977) 39. Chappe, ß., et al.: Naturwissenschaften 66, 522 (1979) 40. Michaelis, W., Albrecht, P.: Naturwissenschaften 66, 420 (1979) 41. Nes, W.R., McKean, M.L.: Biochemistry of Steroids and Other Isoprenoids, Baltimore, London,

Tokyo, University Park Press 1977 42. Eyssen, H.J., et al.: Eur. J. Biochem. 36, 411 (1973) 43. Gagosian, R.ß., et al.: Steroid Transformations in recent marine sediments. In: Advances in Or

ganic Geochemistry, 1979 [Douglas, A.G., Maxwell, J.R. (eds.)], Oxford, New York, Toronto, Sydney, Paris, Frankfurt, Pergamon Press 1980, pp. 407-419

44. GaskeIl, S.J., Egiinton, G.: Nature 254, 209 (1975) 45. Gagosian, R.ß., Smith, S.O.: Nature 277, 287 (1979) 46. Dastillung, M., Albrecht, P.: Nature 269, 678 (1977) 47. Mulheirn, L.J., Ryback, G.: Nature 256, 301 (1975) 48. Ensminger, A., Joly, G., Albrecht, P.: Tetrahedron Lett. 18, 1575 (1978) 49. Schaefle, J., et al.: Tetrahedron Lett. 43, 4163 (1978) 50. Sieskind, 0., Joly, G., Albrecht, P.: Geochim. Cosmochim. Acta 43, 1675 (1979) 51. Shimizu, Y., Alam, M., Kubayashi, A.: J. Am. Chem. Soc. 98, 1059 (1976) 52. Kanazawa, A., Teshima, S., Ando, T.: Comp. Biochem. Physiol. 57B, 317 (1977) 53. Volkman, J.K., Egiinton, G., Corner, E.D.S.: Phytochem. 19, 1809 (1980) 54. Boon, J.J., et al.: Nature 277, 125 (1979) 55. Maxwell, J.R., MacKenzie, A.S., Volkman, J.K.: Nature 286,694 (1980) 56. Seifert, W.K., Moldowan, J.M.: Geochim. Cosmochim. Acta 43, 111 (1979) 57. Conolly, J.D., Overton, K.H.: The Triterpenoids. In: Chemistry ofTerpenes and Terpenoids [New-

man, A.A. (ed.)], London, New York, Academic Press 1972, pp. 207-279 58. Ourisson, G., Albrecht, P., Rohmer, M.: Pure Appl. Chem. 51, 709 (1979) 59. Corbet, ß., Albrecht, P., Ourisson, G.: J. Am. Chem. Soc. 102, 1171 (1980) 60. Spyckerelle, C.: Constituants aromatiques de sediments, Thesis, Strassbourg 1975 61. Wakeharn, S.G., Schaffner, C., Giger, W.: Geochim. Cosmochim. Acta 44, 415 (1980) 62. Spyckerelle, C., et al.: J. Chem. Res. (S), 330 (1977) 63. Albaiges, J., Albrecht, P.: Int. J. Environ. Anal. Chem. 6, 171 (1979) 64. Greiner, A.c., Spyckerelle, C., Albrecht, P.: Tetrahedron 32, 257 (1976) 65. Rohmer, M., Dastillung, M., Ourisson, G.: Naturwissenschaften 67, 456 (1980) 66. Rohmer, M., Ourisson, G.: Tetrahedron Lett. 40, 3641 (1976) 67. LaFlamme, R.E., Hites, R.A.: Geochim. Cosmochim. Acta 42, 289 (1978) 68. Simoneit, ß.R.T.: Geochim. Cosmochim. Acta 41, 463 (1977) 69. Kitadani, M., et al.: Chem. Pharm. Bull. 18,402 (1970) 70. Simoneit, ß.R.T.: Sources of organic matter in oceanic sediments, Thesis, Bristol 1975 71. Boiteau, H.L., Robin, M., Gelot, S.: Arch. Mal. Prof. Med. Trav. Secur. Soc. 33, 261 (1972) 72. Harrington, J.S.: Nature 193, 43 (1962) 73. Harrington, J.S., Commins, B.T.: Chem. Ind. (London) 1964, 1427 74. Geissman, T.A., Sim, K.Y., Murdoch, J.: Experientia 23, 793 (1967) 75. Aizenshtat, Z.: Geochim. Cosmochim. Acta 37, 559 (1973) 76. Niaussat, P., Auger, C.: Compt. Rend. Ser. D. 270, 2702 (1970)


77. Orr, W.L., Grady, J.R.: Geochim. Cosmochim. Acta 31, 1201 (1967) 78. Wakeharn, S.G., et al.: Geochim. Cosmochim. Acta 43, 1141 (1979) 79. Thomas, D.W., Blumer, M.: Science 143, 39 (1964) 80. Giger, W., Blumer, M.: Anal. Chem. 46, 1663 (1974) 81. Blumer, M., Youngblood, W.W.: Science 188, 53 (1974) 82. Youngblood, W.W., Blumer, M.: Geochim. Cosmochim. Acta 39, 1303 (1975) 83. Hites, R.A., Biemann, W.G.: Adv. Chem. Ser. 147, 188 (1975) 84. Hites, R.A., La Flamme, R.E., Farrington, J.W.: Science 198, 829 (1977) 85. Giger, W., Schaffner, C.: Aliphatic, olefinic and aromatic hydrocarbons in recent sediments of a

highly eutrophie lake. In: Advances in Organic Geochemistry [Campos, R., Goni, J. (eds.)], Madrid: Enadimsa, pp. 375-390 (1977)

86. Giger, W., Schaffner, C.: Anal. Chem. 50, 243 (1978) 87. Wakeharn, S.G.: Environ. Sei. Technol. 11,272 (1977) 88. Wakeharn, S.G., Schaffner, c., Giger, W.: Geochim. Cosmochim. Acta 44, 403 (1980) 89. Hase, A., Hites, R.A.: Geochim. Cosmochim. Acta 40, 1141 (1976) 90. Ishiwatari, R., Hanya, T.: Proc. Jpn. Acad. 51, 436 (1975) 91. Brown, F.S., et al.: Geochim. Cosmochim. Acta 36, 1185 (1972)


H.-J. Bolle

Institut für Meteorologie und Geophysik, Universität Innsbruck, A-6020 Innsbruck, Austria

Introduction

The vulnerability of the ecological equilibrium in the biosphere is a matter of growing concern since startling research results about environmental damages due to chemical hazards became evident. These chemical effects are closely interrelated to physical processes in the earth-atmosphere system: The physical properties of the system on the one hand affect transport and reactions of pollutants, while on the other hand chemical constituents in the form of gases or aerosols interact with the radiation regime of the earth due to absorption and scattering processes. More generally we can state that the interdependence of the geochemical system and the terrestrial energetics determine largely the state of the environmental conditions in the biosphere.

It seems appropriate, therefore, to include in a discussion on environmental chemistry a review of the state of our knowledge with respect to the energetics of the earth and its relations to the chemical composition.

The Earth with its atmosphere, oceans, land surfaces and its biosphere is a system with physical and chemical components in which energy of high quality, the solar radiant energy, is transformed into energy of lower quality, the infrared emission to space. Solar radiation is a,high rating energy since it can be utilized for a large number of energy processes, such as: photoionization, photosynthesis, excitation of molecules and atoms, dissociation of molecules, electrical power generation, evaporation, heating and consecutive secondary energy transformations. Infrared radiation is much more restricted in its usefulness since it can exclusively generate heat.

Systems in which a conversion of high rating energy into lower rating energy takes place can be compared with engines which generate mechanical energy fluxes. Following Falk and Ruppel [lJ a system contains an amount of energy which can be exchanged in form of different species of energy. The number of these species depends on the structure of the system, on its "standard

132 H.-J. Bolle

variables". One of the regulating mechanisms which decides upon the species of energy to be choosen in a specific exchange process are the chemical constituents of the system.

Energy changes are described by intensive variables which determine the "level" at which the energy is exchanged, and extensive variables which have the character of quantity. In the earth system the transport of heat plays a dominant röle. The intensive variable of heat is the absolute temperature, T, and the extensive variable associated with heat is called entropy, S. An incremental change of heat, 15QH, can be written 15QH= T dS. Each heat flux, <1>H, is accompanied by a flux of entropy dS/dt (such as any electrical energy flux results from a transport of electric charges, the electrical current): <1>H = T dS/dt for T= const.

The value of the extensive variable is a measure for the quality of the energy. Sensible heat is always transported from higher temperatures to lower temperatures. As all energies it needs a gradient of its intensive variable to be transported. The difference of the entropy fluxes at the source (1) and at the sink (2) of a constant heat flux is given by

(dS\ _ (dS) = 1.- <1>H.2 _ 1.- <1>H. I = Tl - 12 <1>H> 0 , dt/2 dt I 12 Tl 11 12

from which follows (dS/dt}z>(dS/dth; entropy is produced in such a process. In application to the terrestrial system [2] one has to consider that solar radiant

energy is generated in processes which occur at high temperatures (-6,000 K). Thus heat can be extracted from the sun associated with a relatively low entropy. Of this energy a fraction is absorbed by the earth. Here it produces again predominantly heat at relatively high terrestrial temperatures, elose to 300 K. The energy flows to regions where the temperature is lower and ultimately it is emitted as infrared radiation to space which is very cold (- 2 K). This emission takes place at low terrestrial temperatures, near 250 K. The different transport mechanisms involved, conduction, diffusion, eddy diffusion, and large scale eddy transports, have the effect ofheat conduction resistances; otherwise the temperature difference would vanish.

The temperature difference between heat flux source and heat flux sink, according to the above quoted relation, determines the flux of entropy in the system. For a constant heat flux there must be a gain in entropy. The annual average energy flux through the earth system is nearly constant, no substantial energy gain or loss is observed over periods of a year or longer. This on the average constant flux of energy is genera ted at high temperatures and leaves the earth at low temperatures, therefore the terrestrial system produces entropy like an engine, the process is not reversible.

In the course of the involved processes a variety of energy conversions takes place. Radiant energy is converted into chemical and biological energy, sensible heat or latent energy (water phase changes). Supply of liquid and gaseous matter with heat can generate potential energy which may be converted into kinetic energy, and back again to heat due to friction processes before finally the energy is re-ernitted to space.

Radiation and Energy Transport in the Earth Atmosphere System 133

The whole biological cycle is maintained bya very small fraction of the solar radiant energy. Most of its energy is converted in thermodynamical processes which initiate transports in the atmosphere and the oceans at different "scales". The higher frequency (I> l/week) variations are experienced as weather, slower reactions of the system constitute climate. The expression scale is used in connection with weather phenomena to identify specific spatial and/or temporal extensions, such as global, regional, local, annual, seasonal or diurnal events.

Disturbations on aglobaI scale or regionally of the processes by which solar radiant energy is converted into these other energy species may affect the environment considerably and can lead to inadvertent changes with all its socioecological consequences concerning health, food production, fue! consumption, and social behavior. There are the direct impacts on the environmental conditions, e.g. the controlling effect of the ozane layer on the ultraviolet part of the spectrum which is responsible for erythema and skin cancer, or the photochemical reactions which lead to smog if the right trace constituents are present. More important can be indirect changes via the dynamical state of the atmosphere and the oceans, by which climatic zones may be shifted or, more generally, the climate can be affected.

Changes in the Earth's radiative properties can internally be genera ted by a modification of the atmospheric composition, like the increase of the CO2 concentration, the immission of aerosols into the atmosphere, or changes of the reflection properties of the surface due to varying land use methods. The atmosphere-earth system may react to such changes by temperature excursions or by compensating screening effects, like variations in cloud cover. Temperature variations in the atmosphere may in turn affect photochemical reactions, among other pro ces ses. Thus a new photochemical eq uili brium may be established between the reactive components. This new photochemical state could in turn influence the radiation field, a response which is called ajeedback process. We shall see that there exists a number of such feedback mechanisms, which makes a prediction on the ultimate response of the earth-atmosphere system to a disturbation in one ofits components very difficult. Mostly unknown is the reaction of ocean currents, and their possible feedback to the atmosphere.

The further long term development of the Earth's climate may depend to a certain degree on the interplay between atmospheric radiation and chemistry of which a number of aspects are not readily understood. The following secÜons intend to familiarize the reader with the problems involved. Firstly with the interactions between radiation and matter, leading to adescription of the presently applied methods to compute radiation fluxes based upon empirical data. Secondly with the broader aspect of energy transfer in the atmosphere and the oceans, and their relations to the chemical composition of the system. Finally possible climate variations induced by changes in atmospheric or surface quantities and methods to determine critical system parameters will be discussed. The detailed investigati on of this type of problems is the topic of a major research effort during this and probably the next decade, known as the W orld Climate Research Programme. Only partial aspects of this large enterprise can be included here.

134 H.-J. Bolle

Nomenclature, Symbols, and Units

Structure of the Atmosphere and the Oceans

In the atmosphere-ocean system a number of regimes have to be distinguished which differ in their strueture and in which aecordingly different energy transport processes dominate. A schematic overview of the vertical classification is presented in Fig. 1 and the nomenclature used in describing the major features is given below.

The atmosphere is subdivided into the following "spheres". a) The planetary boundary layer (PBL) extends from the surface to about

1 km. H is the layer which is most strongly coupled to the surface. The friction at the ground enforces wind shear and wind veering in the layer. The response of the surface to the incident solar energy determines the temperature structure in this layer. Most of the sensible and latent heat generated at the surface is also released in the PBL. Depending on its temperature and wind strueture this energy as weIl as the pollutants injected from the ground are more or less vigorously transported upward and mixed.

In a very thin surface layer, the laminar sublayer, where wind velocity vanishes, energy is transported by molecular conduction. But very so on, after about 1 mm, it is picked up by the movement of the air which under the influence of solar radiation gets buoyant and transports the atmospheric properties upward by turbulent mixing processes. During the first phase of this upward -transport the fluxes can be regarded as constant. The layer in which these fluxes do not change by more than 10% is the inner boundary layer or Prandl layer. Hs vertical extent depends strongly on the structure of the surface and especially the height of obstacles which cause friction. It may range between a few dm to in the order of 100 m, or about 10% of the whole PBL. Within this inner layer the wind velocity reaches already 70-80% of its frictionless value at the top of the PBL.

The outer PBL is called Ekman layer after the Norwegian scientist [3J who gave the theoretical foundation for the vertical wind profile under the influence of friction. For the mixing and small scale transport of pollutants the wind and temperature profiles in this layer are of eminent importance.

b) The troposphere is the entire layer from the surface to about 12 km and includes the PBL. Hs upper part, the free troposphere, is sometimes decoupled from the PBL by a temperature inversion, whieh hampers exchange between these two regimes. In the free troposphere wind develops according to press ure gradients and is deflected by the Coriolis force caused by the earth rotation. Onee pollutants are mixed high up into the free atmosphere they can be transported over long distanees before they settle down or are washed out by rain.

Throughout the troposphere exists a nearly constant negative mean temperature gradient of approximately r = - 6.5 K km - 1. This is the lapse rate of a weIl mixed convective atmosphere where a temperature distribution results, which enables the air parcels to rise and to sink without exchanging energy with the surrounding air. Such processes are called adiabatic. In actual cases the temperature distribution may depart strongly from this profile. It can be stronger negative, which would favor vertical motion (instable atmosphere), or it can be less negative


1000 km

100 km

10 km

cu 1 km > cu

0 1000 m cu

'" cu > 0 .n 10 m 0

.r; 1m Cl

cu :r:

10 cm

1 cm

r mm

Sea level

-lmm

cu -1 cm > cu

0 -lOcm cu

'" 3 0

cu -1m .n

.r;

0- -10m cu 0

-100 m

-1 km

10km

Mesopause

Stratopause Tropopause

Exosphere

Thermos here

Mesosphere Stratosphere

Free troposphere

Outer boundary layer (Ekman -layerl

Inner boundary layer (Prandl - layei"l

~OgarithmiC wind profile, J nearly constant turbulent momentum and heat fluxes

Mean temperature

Lamminar sublayer (molecular conductionl

Surface interlacial layer 1 o I

1001%

Downwelling solar irradiance in dear water

, I , ,

I , Surlace (wave breakingl layer I

I /

Dep701 euphotic zone ",,//

r,'~;:~layer /"

I Interior

/warm water sphere

layer 'cold water sphere

Bottom water (Deep ocean)

200 250 Temperature in K

Planetary boundary

layer

Troposphere

Height 01 "initial" waves

Warmwater sphere

300

Maximum , wave height

Fig. 1. Vertical structure of the atmosphere and the ocean

135

136 H.-J. Bolle

or even positive which prevents vertical mixing (stable atmosphere). In an adiabatic or neutral atmosphere the pressure reduced potential temperature is

8(z) = T{z} (l05jp{Z})x (1) with

x = 0.287 1, and p{ z} = pressure in Pa 2.

8 is constant for an adiabatic lapse rate. In the stable respectively instable cases it is

d8(z) {>O stable dz < 0 instable .

(2)

c) Separated from the troposphere by a layer of constant temperature, the tropopause, follows the stratosphere which extends from about 12 km (17 km in the tropics, 9 km in the arctic during summer) to 50 km. It is distinguished by its positive temperature gradient up to the stratopause, its upper boundary, where the temperature gradient vanishes again for a few kilometers. The high temperatures at this level ('" 270 K) are due to absorption of solar radiant energy by ozone.

d) The mesosphere is a layer where temperature decends to a minimum of ab out 180 Karound 80 km. It is characterized by intense photochemical processes which lead to a substantial change in atmospheric chemical composition at its higher boundary, the mesopause.

e) The thermosphere which follows on top of the meso pause is heated mainly by oxygen absorption to temperatures up to an average of 1,500 K near 700 km. In the thermosphere further decomposition of the atmosphere OCCurS due to the effect of gravity. At 250--300 km the ratio of atomic oxygen to molecular nitrogen is 1 and these two gases together contribute up to 95% to the total air mixt ure. At 550 km He and 0 are the main constituents mixed at a ratio of 1.

f) Above about 1,200 km, in the exosphere, hydrogen remains as the dominant constituent. In this isothermal layer atoms travel on ballistic trajectories and hydrogen is able to escape to space. The exosphere extends up to about 50,000 km where the magnetosphere starts.

The ocean is confined by two friction layers: the wind forcing at its top and the dissipation layer at the bottom. In both layers exist therefore velocity shear and veering, governed by the same principles as in the atmosphere. There is also some analogy in the thermal behaviour between oceans and the atmosphere. Following Woods [4] the ocean layers can be classificed as folIows:

a) The surface layer is an analogon to the PBL. Wind stress at its surface and response to solar irradiance characterizes its dynamics and its vertical structure. The first 10 m represent the wave breaking layer. Its temperature reflects the diurnal energy transfer processes at the atmosphere-ocean interface. Due to the daily solar irradiance the temperature of the first meter of the ocean can vary in

1 x=RjMA cp , where R= universal gas constant, MA = molar mass of air, cp = specific heat capacity at constant press ure for air

2 105 Pa (Pascal) = 105 Nm - 2 = 1 bar = 1,000 mb = reference air press ure near the surface


still waters by several degrees. During daytime a thermoeline builds up which vanishes at night. Sometimes also a small inversion of a few tenths of a degree may be present very close to the surface, in the interfaeial layer, due to cooling by infrared radiation and evaporation. Differently than for the land surface, the solar radiation is absorbed in a layer of several meters thickness in the oceans rather than directly at the surface. The illuminated layer is called the euphotie zone which can reach down to about 100 m. In very clear waters 1 % of the solar radiation at 425 nm wavelength can reach the 300 m level [122a, 122b]. At higher latitudes a warm surface layer builds up during the summer months with an annual amplitude of the temperature variations of about 10 K. This warm water layer reaches down to about 50 m where a steep thermocline separates it from the cooler deeper ocean waters. In the tropics and subtropics (0-25° latitude) a steep temperature gradient is a permanent phenomenon at depth between 20 and 200 m. The warm surface layer has a depth of about 200 m in the subtropics and 20-50 m in the tropics. Down to this depth, in the eonveetion layer, the ocean is weIl mixed and consequently has rather uniform tem pera tures.

b) Below the steep temperature gradient which separates the upper warm surface layer of the tropical and subtropical oceans from the deeper ocean, the temperature continues to decline, but at a smaller rate, down to 5-10 oe. This main thermoeline reaches down several hundred up to approximately 1,000 m. This region, wh ich is called the warm water sphere, is in exchange with the surface layer due to complicated large scale wind driven circulations (compare Fig. 32). This circulation system extends polewards to about 50° latitude and has its main downwelling branch near 30° latitude.

The temperature continues to decrease below the warm water sphere due to the advection of cold polar waters in the deep ocean. This is the eold water sphere or the deep ocean which extends to approximately 3,000 m.

For the whole layer from about 100 to 3,000 m the name interior layer can be used.

c) Near the sea floor a bottom boundary layer of 10-100 m thickness develops in wh ich the water approach es temperatures of 0 oe in some parts of the oceans. As at the bottom of the atmosphere a viscous sub-layer of 1 mm thickness is present at the ground. But the boundary is not absolutely solid, one has to account for a 10 cm thick sediment percolation layer.

Radiation Tenninology

The basic quantitity to describe the spatial distribution of radiant power in an electromagnetic radiation field is the radiant power area-density per unit solid angle. It is specified as the flux <P of radiant energy (Q) per unit time across unit area into a cone defined by the unit solid angle, and is called radianee with the symbol L. If the radiation field is inhomogenous it is preferable to define the radiance for increments of the area (dA) and of the solid angle (dQ), and to take into account the angle;) between the direction of the electromagnetic wave and the normal of the incremental area dA (compare Fig. 2). The projection of dA in the

138

d'l' { r----I-_

Direction of radiance

c...--L'---r--t-+L.......)---t--__ Nor mal to area

dil = si n~ dß drp

Fig. 2. Radiance geometry and definition of solid angle

direction dQ is cos 9 dA and

d2 cI> L = -co-s-;;:9-;-dA-;--;-dQ=.

The radiance L is measured in W m -2 sr - 1.

H.-J. Bolle

(3)

The radiant flux incident onto unit area from all directions of the hemisphere is called irradiance E. Its unit is W m -2. It can be derived from the radiance by integrating over the hemisphere 2nQo, Qo = unit solid angle,

E= dcI> = 2nQo dA ~ Lcos9dQ W m- 2 • (4)

The radiant energy transported across a defined area during the time t' is measured in W s or Joule (J) and is given by

I'

Q= J cI>dt. (5) o

The energy received during a time interval t1, t2 at unit area is called the radiant exposure H,

12

H= J Edt, (6) 1 1

and is measured in J m - 2 per specified exposure time.


Only for point sources the expression

I=d(/J/dQ (7)

may be used. It is called intensity and is measured in W sr- 1.

If one considers an extended radiation source its radiant power per unit area is expressed by the radiant exitance M, which has the same dimension as the irradiance, W m - 2. The total flux from the source is determined by integration over the surface A of the source:

(/J= J MdA. (8) A

If the source is a black body - defined as an enclosed volume which is in thermodynamic equilibrium at constant temperature- the exitance is conveniently distinguished by an index B:

MB = exitance of a black body . (9)

In practise a black body is constructed as a box with its skin kept at constant temperature. A hole much smaller than the surface of the box serves as emitting area. The radiance resulting from such a source is related to its exitance by (compare Fig. 2)

2n!lo 1t/2

MB= J LB.1-Cos9dQ=2n J LB.1-sin9cos9d9, (10) o 0

where LB • 1- is the black body radiance normal to the radiating aperture and cos9 arises from the orientation of the incremental cone dQ with respect to the normal of the emitting area. For the radiance emerging normally from the black body opening, LB • 1-, the symbol B(T) is introduced. This quantity depends solelyon the black body temperature. The integration of Eq. (10) then results in

M B=nB(T)=nLB.1-' (11)

The radiant energy traversing unit area per unit time, the radiant flux (surface) density is denoted by either M or E.

Spectral radiation quantities are obtained by differentiation of the quantities defined in Eqs. (3) through (11) to one of the spectral parameters: frequency (f or v), wavelength (.,1,) or wavenumber (v). The use of the wavenumber is very convenient for spectroscopic considerations. It is defined by

v = 1/ A. (12)

and is measured in cm - 1. The spectral radiance can e.g. be expressed by

L dL d3 (/J W -2 -1 -1 A= dA. = cos9dAdQdA. m sr nm (13)

if the wavelength is measured in nm, or by

L dL d3 (/J W -2 -1( -1)-1 v= dv = cos9dAdQdv m sr cm (14)

if the wavenumber (in cm - 1) is used.

140 H.-J. Bolle

It should be noted here that the energy in a eertain speetral band }'1" .Al must be the same wh ether A or v is used:

(15)

Therefore the relation holds

(16)

or

(17)

Speetral distributions look therefore quite different if the wavelength or the wavenumber is used as independent variable.

The loss of electromagnetic energy out of a beam either by a change in direction due to scattering or by absorption can be described by Beer's law which states that the loss is proportional to the incident energy:

dL= -aeLds= -Ldu, (18)

where du= ae ds is the differential ofthe optical path length 3, ds is the incremental geometrical path length, and ae is the linear extinction or attenuation coefficient with the dimension m - 1. Alternatively the attenuation can be described by

dL= - ae, mLdm= - Ldu , (19)

where dm = (J ds is the mass sUlface density and a e, m the mass extinction coefficient measured in kg- 1 m2 (or molecules- 1 m2).

For pure scattering and pure absorption separate attenuation coefficients are introduced: the scattering coefficient a (respectively a~ and the absorption coefficient a (respectively a~. Ifboth attenuation processes occur in the same path, the total attenuation coefficient will be

ae=a+a, (20)

and if different (N) substances are involved it is convenient to write

N

du= L (am,i+am,J (Ji ds . (21) i= 1

The reflection by the atmosphere including its aerosols and clouds is essentially a backward scattering process. This is also true for natural earth surfaces where

3 In the u.I.P. Doeument 20 (1978) on Symbols, Units, and Nomenclature in Physies the symbol f1 is preseribed for the attenuation eoeffieient, and the Radiation Commission of the International Assoeiation for Meteorology and Atmospherie Physies (1978) has proposed b for the optieal path length (thickness). Sinee f1 is widely used for the eosine of the zenith angle or the seattering angle in atmospherie radiative transfer theory, the symbol (J is used here for attenuation by seattering and (Je for the total attenuation (extinetion). Other authors prefer the symbols ß or K instead of (Je'

For the optieal path length the symbol u will be used here beeause b may be eonfused with mathematieal symbols. b will exclusively be used here for the optieal depth of the atmosphere, that is the optieal path for a photon travelling vertieally through the atmosphere. Other proposals for symbols made by the Radiation Commission are slightly adjusted in order to eliminate eonfliets with the u.I.P. (1978) doeument


the electromagnetic waves are scattered at soil grains, cell structures in plant leaves or capillary waves on water surfaces. The radiation reflected from natural objects will therefore almost always be a mixture of a specular component and a diffuse component. Some objects like grass may come very dose to a completely diffuse reflector, while quiet water surfaces are almost reflecting like a mirror. In many cases the reflection function (reflected radiance versus direction) can only be obtained empirically. The different reflection geometries are distinguished by the distribution of the incoming and the outgoing radiation. The quantities used for the description of the reflection properties are listed in Table 1. The symbols Qj, Qr stand for the solid angles of the incoming, respectively the outgoing radiation.

The transmitted radiation needs an analogue treatment as presented for the reflectance in Table 1.

If spectral material properties are considered this has to be expressed in terms of a functional dependence like

a(A) = (<PaJ<PO;,.}Je spectral aborptance

Q(},) = (<PeJ<Po;,.}Je spectral reflectance (22)

r(A) = (<PtJ<Po;,.) Je spectral transmittance,

where <P., <Pe' <Pr. and <Po are the absorbed, reflected, transmitted and incident fluxes respective1y, and <P Ä = d<P(A)jdA.

The spectrum whichis of interest for energy processes in the earthatmosphere system can roughly be divided into the following intervals :

A(jJlll) v(cm- 1) Notation < 10- 6 > 1010 y-ray radiation

10- 6 _ 10- 2 106 - 1010 X-ray radiation 0.01- 0.38 26,316 - 106 ultraviolett radiation 0.38- 0.78 12,820 -26,316 visible light 0.78- 2.5 4,000 -12,820 solar infrared radiation 2.5 -1,000 0.1- 4,000 terrestrial infrared radiation

> 1,000 < 0.1 microwave radiation

The spectrum which is of interest for atmospheric research extends over the wide range from about 1 nm to 1 m wavelength as demonstrated in Fig.3. For the energy processes in the lower atmosphere and at the ground only a small fraction of this spectrum is important. 98% of the energy reaching the earth from the sun is concentrated between 0.3 and 4 jJlll and 98% of the longwave radiation emitted from the earth to space is concentrated between 5 and 60 jJlll. Only about 2 octaves of the whole spectrum are responsible for the temperature distribution and biological processes on earth. Shorter wavelengths have special importance for the photo-chemistry in the upper atmosphere because of their high photon energy. Long wavelengths are of interest for radio communication and remote sensing of the earth.

As indicated in Fig. 3 there are different nomendatures in use for the infra red part of the spectrum. For adescription of the energetics of the earth-atmosphere system it is appropriate to discriminate between infrared energy from the sun up

Tab

le 1

. Nom

encl

atur

e fo

r re

flec

tion

qua

ntit

ies

Inco

min

g ra

diat

ion

Dir

ecti

onal

Dir

ecti

onal

Dir

ecti

onal

Hem

isph

eric

al

Hem

isph

eric

al

Hem

isph

eric

al

Out

goin

g ra

diat

ion

Dir

ecti

onal

Dir

ecto

nal

Hem

isph

eric

al

Dir

ecti

onal

Dir

ecti

onal

Hem

isph

eric

al

Ter

min

us

Ref

lect

ion

func

tion

Ref

lect

ion

indi

catr

ix

Hem

isph

eric

al r

efle

ctan

ce f

or d

irec

tion

al i

ncid

ence

Ref

lect

ion

sour

ce f

unct

ion

Dir

ecti

onal

ref

lect

ance

fac

tor

for

hem

isph

eric

al in

cide

nce

Alb

edo

(ref

lect

ance

)

Sym

bol

and

def

init

ion

(Q

Q)=

d

3<t

>,(Q

b Q

,)

y,

" ,

cos9

,·d

Q,·

cos9

jdE

;(Q

JdA

~,

(Qb

Q,)

= n

Qoy

,(Q

j, Q

,)!{2

(QJ

{2(Q

j)=

S y,

(Q;,

Q,)

cos

9,dQ

, 2r

r

J ,(Q

,) =

S

y,(Q

b Q

,) c

os9j

dEj(

QJ

2rr

{2(Q

,)=

J,(Q~;

nQo

{2=<

t>,/<

t>j

Ej=

S

cos9

jdE

;(Q

) W

· m-

2

2rr

Uni

t

sr-

1

Wm

-2

sr-

1

.l:

N :r: ~ co

i2.

"


Fig. 3. Electromagnetic spectrum of the sun and terrestrial sources

N I ~

>-u c

'" :J 0-

'" Li:

'" .CI

10

100

10

E 104 :J c

'" > d 3:

1 m

10 cm

1 nm

lJl

'" > d 3 e ,!,!

E 0

Band designations

p

L

5

C

X

K(Kul

Q(Kal ==0

V

143

Dominating absorption processes

Free electrons

Molecular rotation

Molecular vibrationrotation

Electron excitation

Dissociation

Ionisation

144 H.-J. Bolle

to wavelengths of 2.5 j..Ull (compare e.g. Fig. 14), and the emission of terrestrial sources which just starts to domina te the radiation field beyond 2.5 j..Ull.

A compilation of the symbols used for the different quantities as weIl as of numerical values and units is given in the Appendix.

Elementary Radiation Processes

Relations Between Electromagnetic and Optical Properties of Matter

The radiation which contributes significantly to the energy budget of the earth as a planet consists of electromagnetic waves with wavelengths between in the order of 100 nm and 100 j..Ull. The spectrum in which atmospheric research is interested for different reasons (such as high energy atomic processes or remote sensing) extends further into the shortwave (extrem uv and X-rays) respectively into the microwave region of the spectrum. However, we will restrict ourselves here to the specified part of the spectrum which contains alm ost 100% of the energy. The interaction between the electromagnetic radiation field and matter depends basically on the electrical material constants: specific conductivity, y, permittivity (dielectric constant), e, and magnetic permeability, fl. From these basic material constants the optical refractive index, n, can be deduced which determines the radiative properties of the substance. n generally is a complex number:

(23)

in which the imaginary term nj is responsible for the absorption, the real term nr

for refraction of an electromagnetic wave. If a wave enters an ideally plain surface of a certain material from the vacuum under an angle rt. between the direction of incidence and the normal of the surface, and proceeds within the material under an angle ß between the direction of propagation and the normal, then the real part of the refractive index of the material is given by nr = sinrt.lsinß. It can also be interpreted as ratio of the speed of light in vacuum to the speed oflight in the material.

The absorbing and scattering matter can be treated as a system of damped electric dipols or multipoles which res pond to the exciting electromagnetic wave by resonance at certain frequencies f which are commonly expressed in circular frequencies w=2nfrad S-l. The relation between the index ofrefraction and the electrical material constants can then be written

n2 = fl{e+ i 4:Y}. (24)

The refractive index is thus a function of frequency or of wavelength

Je = elf, c = speed of light, (25)

and can as weIl depend on temperature and pressure. All optical phenomena can in principle be deduced from a basic set of electro

magnetic field equations by application of Maxwell's theory [5, 6]. The basic


material constants and resonance frequencies have to be known, for instance from solid state physics, in order to obtain results which describe the observable phenomena. The optical properties of spherical particles suspended in the air, e.g., can be computed by Mie's theory [6, 7] which is entirely based upon Maxwell's equations and continuity requirements. If the scattering centers are small compared with the wavelength, Lord Rayleigh's theory [8a-c] applies which is a special case of the more general theory of Mie. Progress has also been made during the last decennia in treating other regularely shaped particles in this way [9, 10]. In practice the application of the theory leads to numerical difficulties if chemically inhomogenous materials with irregular surfaces have to be considered. Thus aerosol particles with irregular shapes, as weIl as the natural earth surface, can only be treated by means of empirically determined optical "bulk" parameters. The expression "bulk parameter" suggests that macroscopic average parameters of certain ensembles and not of individual elementary particles are measured.

In the following treatment only those properties will be discussed which are necessary for an understanding of the energy processes connected with radiative transfer. Optical phenomena resulting from refraction and polarization will not be considered. The interested reader is referred to more elaborate treatments of these subjects [11-16].

Molecular Scattering

For the wavelength range in which most of the radiative energy is transported in the earth-atmosphere system (). > 100 nm), molecules can be treated as infinitely small spheres in which charge displacements are induced by an external electromagnetic field. If the external field is genera ted by a polarized sine wave, then the charge displacements are periodically altered and the model of an oscillating dipole can be applied. Such an oscillating dipole generates a secondary wave like a Hertzian dipole antenna of macroscopic dimensions. The irradiance produced by this secondary wave in large distances from the dipole was derived by Hertz [17]. The radiant energy at wavelength ). can be represented by the Poynting vector S averaged in time and over an incremental volume 4 which can be identified with the radiant flux surface density M at the distance r of the dipole.

The mathematical relation between these quantities is given in the next section, Eqs. (35) and (36). Once this relation between the electromagnetic properties of the material and the exciting electromagnetic field is established, it is relatively easy to deduce an equation between the refractive index of the material and the macroscopic attenuation coefficient as defined by Eqs. (18) or (19). For scattering centers which are much smaller then the wavelength, as is the case for scattering molecules, the relative simple relation

( ')_ 32n3(n-1)2 -1 aRA - 3N).4 m (26)

4 The Poynting vector is originally defined as energy l1ux per unit area through the surface of a volume, where only its normal component needs to be defined, and is identical with the vector product of the electrical field strength vector E and the magnetic field strength vector, H

146 H.-J. Bolle

is obtained. The index R denotes that the attenuation eoeffieient is related to moleeular or Rayleigh scattering N is the number of moleeules per unit volume [see Eq. (41)], n the index ofrefraetion [Eq. (30)].

The radiation is not seattered isotropieally into all direetions, but has a eharaeteristie angular distribution. As will be shown in the next seetion, the relative distribution (normalized to one if integrated over all angles) for Rayleigh seattering ean be expressed by a scattering phase function

(27)

where 9 is the seattering angle (eompare Fig. 4). The two terms at the r.h.s. result from the two polarized eomponents of natural light. It ean be shown that the eos29 term results from the eleetrie veetor parallel to the plane defined by the ineident beam and the line of sight of the observer. It is responsible for the anisotropie seattering eomponent. In all direetions exeept for 9= 0, 90, and 180° exists a mixture of polarized and non-polarized light. The degree of polarization ean be expressed [eompare Eqs. (44) through (48)] by:

p_ h(9)-1 jj (9) _ 1-eos2 9 - 11.(9)+111 (9) - 1+eos2 9 (28)

(29)

whieh indieates a eomplete polarization (P= 1) at 9= 90°. In the sky the polarization is indeed maximum near 90° off the direetion of the sun, however it beeomes never 100%. This observation ean be explained by multiple seattering in the atmosphere-earth surfaee system, and by the depolarizing effeet of aerosol seattering. 5

The Rayleigh seattering eoeffieient depends explieitely on the fourth power of the wavelength. But also the refraetive index of air depends slightly on the wavelength. Its empirieal numbers ean be approximated by the following formula (at 288 K and 1,013.25 mb, A in j.Ull):

( -1).106 =64328 29,498.10 255.40 n . +146-A 2 +41-A 2 (30)

or with less aeeuraey by

(n-1) 106 ~N 00(1 + 7.52·10- 3 A -2), (31)

where A< 20 j.Ull and

N 00 = (n oo -1) = 77.6 p/T. (32)

noo is the refraetive index for wavelengths mueh larger than 20 j.Ull or v ~ 500 cm - 1,

P is the air pressure in mb and T the temperature in K [18]. The wavelength IS In j.Ull.

5 Three neutral points are observed in the sky: The point of Arago, 15-20° above the anti solar point and the points of Babinet and Brewster 15-20° above respectively below the sun


The total spectral dependence of the Rayleigh scattering coefficient can be approximated by

(33)

The A - 4.09 dependence of the scattering coefficient explaines the blue color of the sky: the short wavelengths are scattered much stronger than the longer ones. This blue color would, however, not occur, if in the atmosphere the distances between the molecules would be constant like in a crystal or a liquid. In that case the light scattered from the different centers would be coherent with fixed phase differences, and the scattered light would be extinguished due to interference. The Brownian movement of the molecules destroys this coherency in gases, and the intensities rather than the amplitudes of the scattered light add. This atmospheric property has been used in the deduction of Eq. (26) respectively (42).

The treatment of the scattering centers as small particles holds up to radii of about r<0.03 A within 1 %. Precise measurements have indicated that molecules can not be regarded as ideal spheres. Some depolarization is always present due to the anisotropy of the molecules. In order to account for this effect an anisotropy factor (6 + 3 pJ/(6 - 7 pJ ~ 1.06 is applied as a correction factor to the extinction coefficients. Pn is defined by the residual depolarization at 90° with respect to the direction of the incident light:

(34)

Deduction of the Rayleigh Scattering Coefficient and Phase Function

This section gives the mathematical background for the conclusions drawn in the previous one. It gives an instructive example how radiation quantitities can be deduced from basic material constants.

The theoretical treatment ofthe Rayleigh scattering starts with the formula ofthe radiating dipole derived by Hertz [17], which relates the Poynting Vector [19, 20] of the radiant l1ux density to the material properties expressed by the resonance frequency and the electrical vector of the exciting field:

(35)

where

M A = ~~ = spectral radiant l1ux (area) density in W m - 2 m - 1

C =velocity oflight (2.9979.108 m S-I)

Bo = dielectricconstantorpermittivityofvacuum= 8.854.10- 12 Fm -1 (1 Fm -1 = 1 A· sjV . m) A =wavelength e =elementary charge (electron charge = 1.602.10- 19 A s) m =mass of charge (electron mass = 9.1095 . 10- 31 kg) Wo = (kjm) 1/2 re sonant circular frequency (k = restoring force per unit displacement) in s - 1

w= 2ncjA =electromagnetic circular frequency in S-1

Eo = maximum value of external electric field strength vector in V m - 1

!p = angle between direction of observation and the direction of the dipole axis = distance dipole-observer.

For microscopic dimensions the term e2 jm (w5 - w2 ) depends on the polarizability ofthe radiating gas moleeule, and this quantity can be related by a formula derived by Lorentz and Lorenz [21, 22] to the refractive index n and the number N of dipole oscillators per unit volume:

(36)

148 HA. Bolle

The intensity as defined by Eq. (7) can be expressed by I=d<1>jdQ=Mr2 using M=E as defined in Eq. (4) for the flux density and for the secondary wave per molecule it results with Eqs. (35) and (36):

(37)

where (38)

is the averaged value in W m - 2 of the Poynting vector of the incident wave polarized normal to the direction of propagation. The ratio of the intensity scattered in the direction cp to the incident irradiance is the directional scattering cross section ö'(cp). For Rayleigh scattering

(39)

With n ~ 1, and consequently n2 + 2 ~3, (n 2 _1)2 = [(n + 1) (n _1)]2 ~4(n _1)2, it follows from Eqs. (37) and (39) at wa velength A:

(40)

The integration of the first expression in Eq. (40) over all angles cp yields the total scattering cross section in m2 for one molecule or a small spherical particIe:

(41)

In a unit volume there are N scattering centers. Therefore the Rayleigh scattering per unit volume and unit solid angle can be described by

(42)

or with Eq. (40) by

(43)

If the preposition is dropped now that the primary wave is polarized and natural light is considered then the primary field can be regarded as composed of two linearly polarized components, one parallel to a reference plane (MII) and one perpendicular (M.c) to it. The reference plane can arbitrarely be choosen. If M is considered as the total radiant flux density, then the two polarised components each carry half of the total radiant flux:

MJ.=M II =Mj2. (44)

The intensity of the unpolarized light can then be expressed according to Eqs. (39) and (40) by

I.(CPl' C(2) = Iu(cpr) + III.(CP2) = UR(Ä; CPl) M u + aR(}'; C(2) M II,,,

_ M" n 2(n 2 _1)2 ( . 2 . 2 ) - T N2J..4 sm CPl +sm CP2 . (45)

Here cP 1 and CP2 are the angles between the direction of observation and the direction of the electric vector of the two polarized incident waves respectively. The angle between the direction into which the incident waves propagate and the direction of observation, the scattering angle 9, can be expressed by CPl and CP2 due to spherical geometry in the following way (compare Fig. 4):

or

By substitution of (47) into (45) and multiplication with N it results in analogy to Eq. (43):

n2(n 2 _ 1)2 YR(J..;9)= 2NJ..4 (l+cos2 9)m-1sr- 1,

(46)

(47)

(48)


Fig. 4. Geometry for scattering

. _ 190 ) - cos 'P, sm ~ - cos -}! - cosl90- \1')

cos l-' _ cos 11'2 - cosI90-0) _ cos29', + cos 2rp2 - sin 2 z?

149

Direction 01

propagation

which is the Rayleigh scattering function expressed in terms of the scattering angle:} for unpolarized radiation. By integration over all scattering angles the total valurne scattering caefficient due to Rayleigh scattering can be obtained,

41t!10 1t

O"R(A)= J YR(A; 9) dQ=2n JyR(A; 9) sin9d9, o 0

which is equivalent to the linear attenuation coefficient due to Rayleigh scattering. rr

Since J (1 + cos 2 ep) sinepdep = ~ it follows o 8n3(n2 _1)2

O"R(A) = 3N A4

(A) "" 32n3(n _1)2 O"R - 3NA4

or

The distribution of the scattered intensity is generally described by the scattering phase functian

4n p(9) = -y(9)

0"

(49)

(50)

(26)

(51)

which is normalized to 1 in the conservative (= non absorbing) case. y is the directional volume scattering cross section [defined in Eq. (42)]. For Rayleigh scattering the phase function is accordingly [Eq. (48)]

(27)

The intensity scattered into different directions depends only on the scattering angle 9 and not on any choosen reference plane or the orientation of the electrical vectors with respect to this plane. We can therefore choose the plane defined by the direction of observation and of the incident light as reference plane with one electric vector vertical, the other parallel to this plane. Then the vertical vector gives

150 H.-J. Bolle

rise to a scattered intensity according to Eq. (40) setting 'P = 'PI = 90° with respect to the reference plane. This component is therefore taken care of by the 1 in the parenthesis of the r.h.s. of Eq. (27).

Aerosol Scattering

If the radius r of the scattering centers becomes larger than 0.03 A the Rayleigh approximation of scattering theory becomes invalid. The bulk of atmospheric aerosols has indeed radii comparable with the wavelength of the scattered solar radiation.

The scattering by particles of sizes comparable with the wavelength has first been treated theoretically by Mie [7]. The leading idea of the Mie theory is to treat the aerosols as isotropie, spherical particles with an uniform complex index of refraction. The incident wave is able to excite all kinds of electric and magnetic oscillations inside the particle and accordingly radiation from multipoles of different orders. The secondary waves from these multipoles combine and interfere, and at the skin of the particle the condition of continuity must be observed with respect to the external field. The condition is that the tangential component of the electric and magnetic field vectors of the secondary waves and of the incident wave must be continuous across the surface. Essentially from this boundary condition the governing equations for the structure of the scattered field outside the particle can be deduced mathematically.

Rather than to develop the complicated theory (see e.g. [6]), only some results and the differences to the Rayleigh case will be referred to.

In order to derive formulations of general character for the scattering by aerosols of different sizes the dimensionless size parameter IY. is conveniently introduced, which is the ratio of the particle radius to the wavelength multiplied by 2n:

1Y.=2nr/A. (52)

Extinction and Scattering Cross Sections of Monodispersions

The total Mie extinction cross seetion is described in terms of an efficiency factor or Mie coefficient Qeff which is the ratio of the Mie extinction cross section, UM, defined analogue Eqs. (39) respectively (41), to the geometrical cross seetion nr2 :

(53)

where the wavelength dependence is not explicitly no ted. It is assumed that the scattering is rotationally symmetrie with respect to the

scattering angle ;) around the direction of propagation of the incident wave (compare Fig. 4).

In analogy to Eq. (20) the efficiency factor Qeff can be splitted into two terms:

(54)

where Qs is the scattering efficiency factor and Qa the absorption efficiency factor. For very small particles Qeff is much less than unity: the scattering efficiency

of the particles is much smaller than their geometrical size would require.


It follows a steep increase with (X until a maximum is reached near (X = 6. The exact position and the maximum value depends on the index of refraction respectively the chemical composition of the aerosol. For water droplets (n r = 1.33) the efficiency factor rises to about 4 around (X=6. For larger (X the scattering efficiency starts to oscillate with decreasing amplitudes around the value of 2 at which it finally converges for (x--+OO in agreement with the theory of geometrical optics. With increasing absorption (= increasing imaginary part of refractive index) the amplitude of the oszillation is damped. At (X 4! 1 the Rayleigh approximation holds with a wavelength dependence A -4 while for (X--+ 00 the wavelength exponent becomes zero. In between these extremes all values between - 4 and '" 1. 7 occur for the exponent of the wavelength dependence.

Angular Distribution of Scattered Intensity

The directional scattering cross section, aM,s(8), which is related to Qs [Eq. (54)] as aM(8) is to Qeff [Eq. (53)], is a measure for the angular distribution of the scattered radiation. It is closely related to the phase function [Eq. (51)], which defines the normalized directional scattering of a unit volume while the cross section is related to one particle.

In contrary to the Rayleigh scattering the Mie scattering is distinguished by a marked foreward intensity. With growing particles more and more radiant energy is scattered into a narrow cone in the direction of the incident primary wave. The scattered radiance drops by orders of magnitude in approaching 90° offthis direction and oscillates in the backward direction. The angular dependence between 90° and 180° can be very complicated for distinct size classes as a result of the interference of the partial waves generated at different parts of the aerosol particle. This asymmetry of the scattering indicatrix or phase fonction p(8) first grows with increasing size parameter (X. For e.g. (X= 5 the intensity for scattering angles > 60° is only a few %0 of that in foreward direction. For large size parameters especially the polarized components start to show marked minima. This structure is somewhat smoothed in the total scattered intensity.

A convenient measure for the asymmetry of the phase function with respect to the forward-to-backward scattering ratio is the asymmetryfactor defined by the phase function weighted average of the cosine of the foreward scattering angle:

+1

g=<cos8)=i J p(8)cos8d(cos8). -1

(55)

Up to (X = 0.5 the scattered radiation is still nearly completely polarized at 8=90°. The polarization decreases, however, as the size parameter grows and the location of maximum polarization is shifted towards larger angles.

Absorption in the particles does not only affect the total scattered intensity but also the shape of the phase function. For e.g. (X=5 there is still not much difference to the nonabsorbing case as long as nj ~ 10 - 2. However, with increasing nj the minima sharpen for the component with the electric vector perpendicular to the plane of observation until they vanish at about nj= 1.0. For the parallel component oszillations in the angular energy distribution start at about nj=O.l and increase with the value of nj (see e.g. [11, 23J).

152 H.-J. Bolle

Albedo of Single Scattering w A quantity which can be used as a measure for the absorption in the aerosols is the albedo for a single scattering process

- ! (Je- a 1 -())=(J (J = --= -0( e (Je

(56)

which is the ratio of the flux scattered in a volume to the total scattered and absorbed flux. fi.=a!(Je is the absorption number.

Scattering in a Hazy Atmosphere

Under natural conditions in the atmosphere one never has to deal with monodisperse aerosols but with a more or less broad range of sizes. One common approach to represent the aerosol size distribution follows the proposal of Junge [24] in which the number of particles per unit radius interval and per unit volume is given by a logarithmic function of the radius. But there are also Gaussian size distributions in use [25a-d, 26], compare the next section.

The wavelength dependence of the extinction coefficient results from the contributions of all monodispersions [27]. For natural aerosols the superposition of the individual scattering coefficients produce a mean overall slope of the extinction coefficient of approximately A -1.3 (the exponent can range between 0.12 and 2.3). It is clear that at this point experimental verification is necessary. Often the chemical properties of the aerosol are not well enough known to compute the scattering properties of the mixture, which makes it necessary to determine them by direGt measurements.

Representation of Aerosol Size Distributions

If n(r)/r denotes the number of particles per unit volume and per unit radius in the size range r, ... ,r+dr (r=radius), then the total number density of particles between the lower limit of the distribution, rl, and r2 > rl is

or

N(r)= 'f n(r) dr " r

dN n(r)=-dl . ogr

Between about r= 10- 1 /lffi and r= 10 /lffi n(r) can be approximated by apower law function

or, with dlogr=0.434dlnr, dlnr=dr/r;

dN/dr=c'r- v.- 1, with c'=0.434 c,

from which follows '2

N=c' J r-(vo+l)dr.

"

(57)

(58)

(59)

(60)

The constant c depends on the concentration and v* determines the slope of the distribution curve. Over a wide range of the distribution (0.1 < r < 1 /lffi) a value of v* = 3 can generally be accepted.


The monodisperse volume extinction coefficient can according to Eq. (53) and in analogy to Eqs. (42) and (49) be represented by the extinction efficiency Qorr:

(61)

For a broader distribution with particles from ri to r,+ 1 it has to be summed over all r in this interval. With Eq. (60):

'i+ 1

O"M,.(rj,ri+l)=nc' J r-(VO+l)r2Q.rrtr)dr, (62)

"

(63)

are introduced:

(64)

K i can be computed for a number of narrow intervals from which the total extinction coefficient is determined by summation over the intervals

(65)

The power law distribution, Eq. (58), represents the observed average natural tropospheric aerosol size distribution quite weIl. For special cases a modified gamma function or a superposition of two or more of such functions may improve the description of the actual size distribution. In the modified gamma representation the size distribution is given by an expression of the type

n(r) = ArPe-B,' , (66)

where A, B, ß, and )' are empirical constants, The maximum of the distribution occurs at

rmax=V ßIB)' (66a)

and the total concentration is given by P+l

N= AB~ y r(ß;I), (66b)

where r is the gamma-function. The slope of the distribution in a double-logarithmic plot is given by

dlogN(r) = -ß[1-(!....)']= -(v*+I). dlogr r m

(67)

Here (v* + 1) is the equivalent slope of the power law distribution, Equation (67) allows a conversion between the two representations Eqs. (58) and (66).

Absorption

Radiant energy is absorbed by a number of processes which reach from a successive damping ofthe electromagnetic wave by multiple reflection in cavities - as they exist in certain aerosols like soot and at the solid surface of the earth - to the highly energetic ionization of atoms. In solid bodies radiative energy is exciting vibrations of lattice structures or single molecules. The primary vibrations are immediately damped by the inter-molecular forces and the radiant energy is distributed within the absorbing substance as heat, giving raise to a general increase of molecular movements respectively temperature. Certain materials,

154 H.-J. Bolle

such as amorphous carbon, are able to absorb a very broad band of electromagnetic energy, other materials with distinct lattice structures absorb selectively at wavelengths where the characteristic frequencies are in resonance with the frequencies ofthe incident waves. Due to defaults in the lattice structure they are, however, often also absorbing radiation in more broader spectral bands.

It normally does not need much energy to excite the rotation of agas molecule whieh is not imbedded in a lattice structure. Long wavelengths (A. > 20 1Jlll) can therefore be absorbed by exciting rotationallevels which correspond to the photon energy hcv of the incident radiation.

The excitation of vibrationallevels requires more energy and accordingly takes place at shorter wavelengths than the excitation of rotational levels, normally from the near infrared ( '" 1 1Jlll) to ab out 20 lJlll wavelengths. The excitation of a vibrational level is accompanied by the excitation of a rotational level so that these absorptions exhibit a complicated band structure.

The absorption process requires that the magnitude of an existing dipole moment or of a higher moment is changed. Only those molecules therefore absorb photons whieh have a permanent dipole (or multipole) moment. Molecules of two identical atoms like O2, N2, H2 do not normally have this capability because their symmetrie structure prevents a charge separation or polarization within the molecule. Only in case of deformation due to pressure or strong external electrical or magnetieal fields polarization may be induced and excited levels appear. Multipole transitions occur with much smaller probability than dipole transitions.

Still more energy is needed to induce changes in the electronie states of molecules or atoms, some of which occur in the visible part of the spectrum (0.30< A. < 0.72 1Jlll) but most ofthem in the ultraviolet(A. < 0.38 J.UIl}. The excitation of moleeules can finally result in a dissociation of the compound into two atoms or one atom and a smaller molecule or radieal. Absorption of energetic photons mayaiso lead to aseparation of one or more electrons from a moleeule or atom, leaving back an ion. This reaction is called photo-ionization.

Absorption by Aerosols

The absorption number defined by Eq. (56) represents the absorbed fraction of electromagnetic energy at a single scattering process. This quantity can be computed from Mie theory if the refractive index is known and if the particle is spherical or elliptieal and has a simple symmetrie structure. It is much more difficult to determine this quantity in mixtures of aerosols as they exist in nature, and it is even more difficult to determine the absorption of an aerosol layer because here the multiple scattering between the different scattering centers has to be accounted for. An ensemble of suspended aerosols generally has optical properties that differ from those computed from the refractive indices of its individual chemical components for single scattering processes.

The application of different measuring techniques [28, 29, 30a, 30b] and of theoretical deductions [31-33] have resulted in improved insights into the absorption properties of aerosols [34]. The single scattering albedo at 0.55 lJlll wavelength, whieh is often taken as reference for the visible part of the spectrum,


ranges between 0.647 for industrial haze and 1.0 for maritime and stratospheric aerosol. Soot has a value of 0.209. At 10 ~ the average extreme va lues are 0.173 for urban-industrial aerosol, 0.692 for maritime aerosol and 0.0 for soot [35].

Line Absorption

Absorption due to a transition between energy levels of gas molecules occurs in narrow spectral intervals defined by the width of the energy levels between which the transition occurs. The line absorption coefficient al is defined by an intensity function S6, the line strength, and a shape-factor which depends on the kind of process wh ich broadens the energy level. According to the Heisenberg uncertainty relation each level has a natural width which is so small that it needs not to be considered in practical applications. In denser atmospheres broadening due to molecular collisions is the dominant factor for the width of the energy levels. The uncertainty of the energy levels is reflected in the half width (I. of the slope of the absorption coefficient which can in the vicinity of the line center 7 be described by a Lorentz line shape around the central wavenumber vo:

S (I. adv)= n(v-vof+(l.2· (68)

The line width ((I.) decreases proportional to the pressure. In the upper atmosphere, at lower pressures, therefore another process becomes important: the Doppler li ne broadening due to the radial velocity of the molecules with respect to the observer (Fig. 5).

Spectra of atIhospheric molecules have in detail be treated by Herzberg [37], Penner [38], and Goody [39].

The Lorentz half width due to molecular collisions, Cl, is a function of press ure and temperature and can be represented by

Cl=ClO(~) (Ta)", Pe,O T

(69)

where e,g. for H 2 0 -0.833<n<0.045 (mean value n= -0.5). The equivalent press ure is Pe= p+ (B-l)Pi, where P= total pressure, Pi=partial pressure of absorbing gas i, B= self-broadening coefficient. Clo = Cl(Po, Ta) is in the order of 0.1 cm - 1 (maximum ~0.11 for 0 3, minimum 0.0086 for high rotational transitions in water vapor). Po= 105 Pa, T=296 K, and B has a value of ",,5.

As the density of the atmosphere drops, the Doppler effect resulting from the movement of the molecules with respect to the observer starts to ga in in importance until it dominates the shape of the line. The general formula for a line with pressure and Doppler broadening is called a Voigt profile [38, 40]:

[ In2 _, - )2] __ SClL~n2+00exp- ClÖ(V-VO "

av(v)- -- - S ( ')2 2 d\, (70) nClD n -00 v-v +ClL

where

_ vo~RT ClD(V)= - --ln2 c M

(71)

R = universal gas constant, M = molar mass, c = velocity of light.

6 This quantity has nothing in common with the Poynting vector for which the same symbol was used 7 The Lorentz li ne shape does not hold beyond some halfwidths off the center, where often a steeper

decrease is observed

156

NE u

'0>

1:.-0

H.-J. Bolle

6000

180

5000

4000

3000

0.1 0.2

0~--~~~~~~==~~3~~~~~~~--~10~-2--5·10-

V-Va (ern-I)

Fig. 5. Line mass absorption coefficient a" in units g-I cm2 computed for the water vapor line 66-50

at 18.477 J.lm. a: Doppler, b: Lorentz, c: Voigt profile for 30 km altitude. Insert: Lorentz line profile for 3.57, 8, and 12 km altitude. After Bolle et al. [36]

In Fig. 5 the absorption coefficient has been computed for 30 km height with the Doppler, the Lorentz and the Voigt line profile wh ich demonstrates the differences in shape produced by the three mechanisms of line excitation. The integral over the line must per definition be equal in all three cases, since

s= J alv) dv (72)

independently of I = L, D or V. The temperature dependence of S is related to the population of the energy levels and can be described by the energy of the level to which the transition occurs [38,41].

The amount of energy absorbed in a line can be found in the following way. Given an optical path

(73)

where amI is the mass absorption coefficient for a line. The transmittance through the layer can be computed due to application of Beer's law from

(74)


100 mb, 217°K

-1 10 mb, 228°K 1 mb, 271 0 K

T 0.1 mb, 237°K

E 0.01 mb, 181 °K

~ -2

z; "0 -3 3!

C -4 ~

0 > ::J rr -5 Q)

01 S!

-6

-11 -10 -9 -8 .7 -6 -5 -4 -3 -1 o log Water vapor mass (gcm-2)

Fig. 6. Equivalent width of a strong rotational water vapor line near 50 jllIl wave1ength computed for live different levels of the atmosphere [42]

The difference 1- !(v) = lJ((v) is the absorptance of the layer and

(75)

is called the equivalent width which defines a rectangular spectral interval of zero transmission circumscribing the same area as the line. As long as the line center is not saturated the equivalent width grows linearly with the mass in the optical path. Then the center of the line becomes saturated and there is a transition region in wh ich the absorption changes less with growing mass. If in a central portion the line is completely opaque radiant energy can only be absorbed in the wings. In this case the equivalent width grows proportional to the square root of the mass:

W=2~ (76)

which is called the strong line approximation. The three regions can clearly be distinguished in Fig. 6.

Band Absorption

In absorption bands the situation is complicated by the overlapping of the lines as they grow with increasing absorber mass. A thorough treatment of radiative transfer in bands can therefore only be achieved by line-by-line calculations. According to the magnitude of their contributions, a large number of lines is normally included by summing up their individual contributions to the optical path (N =number of contributing lines):

N

u(v) = J I amli(V) dm. i= 1

(77)

Fortunately the lines are known for most of the important atmospheric species with reasonable accuracy [43--45].

z o ;n '" i '" z ... e:

158 H.-J. Bolle

Band absorptances are also known from extended laboratory studies [46--49] from which approximative formulas of the type

S ct(v) dv= c[mp!]b (78) or

S ct(v) dv= C+D logw+ K logp (79)

have been derived. The coefficients a, b, c respectively C, D, and K vary from band to band. b is always close to 0.5. Pe= p+ Pi is again the equivalent pressure, P= total pressure, Pi = partial pressure of the investigated gas. Numbers of the constants are listed in Table 2c.

Rotation Spectra

üf the rotation spectra only that of water vapor is of significant importance for the transfer of radiant energy in the atmosphere. It extends from approximately 10).lIIl to the microwave region and is centered at 202 cm - 1 or ab out 50).lIIl. The water vapor moleeule is an aspherical top, and because of its strong asymmetry and large permanent dipole moment a wide range of rotationallevels is excited. The structure of the band is highly complicated and exposes an alm ost statisticalline distribution. A small section of this band is shown in Fig. 7 which demonstrates the high opaqueness of the atmosphere due to this band beyond 15 ).lIIl wavelength.

11111111111 1111 1

~ 1.0

0.5

550 560 570 580 590 600 6'0 620

WAVENUMBER - WELLENZAHL "' ,m-I

Fig. 7. Computed atmospheric radiance to space between 550 and 630 nm. The computation shows a superposition of water vapor, carbon dioxide and nitrous dioxide lines. The regular nitrous dioxide band is shown at the top in transmission in contrast to the highly irregular water vapor band

630


Vibration-Rotation Spectra

For atmospheric gases vibrational transitions occur of wavelength between about 20 llIll and lllIll. Simultanously with the vibration also rotational levels are excited. Only transitions are allowed which are accompanied by a change of the electric dipole moment or, with much less probability, of higher moments (magnetic dipole, quadrupole etc.). The possible transitions can be computed from quantum theory. They are restricted by selection rules which take care that certain symmetry conditions are not viola ted.

The energy involved in the rotational transition can either be added to or subtracted from the vibrational energy. A vibrational band therefore consists always of two "branches", one high-frequency wing where the rotational energy is added to the vibrational energy (R-branch) and one where the rotational energy is subtracted (P-branch) which is at the low frequency side of the vibrational transition. For multi-atomic moleeules sometimes also transitions are allowed which are not accompanied by a change of the angular momentum. In these cases a third branch appears right at the center of the vibration band which is normally very narrow and can hardly be resolved into lines. It is called Q-branch and contains lines resulting from transitions of corresponding rotational levels at the two involved vibrational states.

Diatomic moleeules and linear symmetrie multi-atomic moleeules generate vibration-rotation bands of nearly equally spaced rotationallines (compare Fig. 7). Asymmetrie moleeules like water vapor and ozone reflect in their vibrationrotation spectra the complicated structure of the rotational band. They appear therefore to be much more irregular (compare Fig. 16).

The irregular appearance ofthe spectrum is sometimes considerably increased by the superposition of isotopic bands with slightly shifted frequencies and high er order combination bands. Within the region of the 15 llIll COz band (605~750cm-l), e.g., 83 bands have been identified, which are partly very weak and do not much affect the absorption at small optical pathes in the region of the dominating fundamental COz vibration-rotation band at 667.379 cm- 1 (14.981lIll). But some of them become very important at high er pressures and elevated temperatures ("hot bands").

The most important atmospheric species with their bands are listed in Table 2 a, and bands of minor atmospheric constituents in Table 2 b.

Quasi-Continuum Water Vapor Absorption

In between the major far infra red water vapor bands centered at 2.7 1lIll, 6.3 1lIll, and 50 llIll an absorption has been measured that still is not completely explained. It is very likely that part of it results from the accumulated effect of the wings of far distant strong absorption lines. Another explanation is that HzO dimers [(HzO)n molecules] which form in the atmosphere exhibit absorption bands just in the "window" regions. No re solvable band structure of this attenuation has yet been found and since at least part of it is very probably caused by far distant line wings, it is sometimes referred to as "quasi"-continuum.

Since no quantitative explanation can be offered for this phenomenon it is also not possible to postulate apriori laws for its pressure and temperature

160 H.-J. Bolle

Table 2a. Major vibration-rotation bands of atmospheric gases [43]

Gas Spectral Range Number Dominating band ern-I llII1 of

bands Identification Center wavenumber

H20 1,300- 2,400 4.2 - 7.8 7 010- 000 1,594.736 2,550- 2,800 3.6 - 4.1 3 3,000- 4,500 2.2 - 3.3 15 100- 000 3,657.054

001- 000 3,755.924 4,500- 6,200 1.6 - 2.2 11 011- 000 5,331.245

101- 000 7,249.93 6,200- 8,000 1.25- 1.05 8,000-11,000 0.9 - 1.0 13 111- 000 8,807.00

CO2 450- 800 12.5 -22 125 01101--0000 I 667.379 800- 1,200 8.3 -12.5 26 00011-10001 960.959

00011-10002 1,063.734 1,200- 2,000 5 - 8.3 24 2,000- 3,000 3.3 - 5 120 000 11-0000 I 2,349.146 3,000- 4,000 2.5 - 3.3 95 10012--00001 3,612.844

10011--0000 1 3,714.781 4,000-10,000 1 - 2.5 149 20013--00001 4,853.62

20012-00001 4,977.830 200II-DOOOl 5,099.66

0 3 700.930 14.267 I 000- 000 1,000- 1,200 8.3 -10 7 001- 000 1,042.096 2,110.79 4.7376 101- 000 2,785.241 3.5904 111- 000 3,041.200 3.2882 003- 000

N20 500- 600 16.5 -20 11 01 10- 00°0 588.767 696.140 14.365 1 938.849 10.651 1

1,100- 1,400 7 - 9 20 10°0- 00°0 1,284.907 1,400- 2,000 5 - 7 6 2,000- 2,300 4.3 - 5 18 00°1- 00°0 2,223.756

01 1 1- 01 ' 0 2,209.523 2,300- 5,000 2 - 4.3 49 2000- 00°0 2,563.341

10°1- 00°0 3,480.821

CO 2,169.836 4.6086 1- 0 2- 0 3- 0

CH4 1,297.88 7.7049 V4

1,305.914 7.6575 V4 1,305.914 1,533.289 6.5219 V2

2,600 3.846 2V4

2,818 3.549 } V4+ V2 2,838 3.524 3,009.53 3.3228 V3

3,018.920 3.3124 V3

O 2 5,241.09 1.908 a-X 6,325.99 1.581 7,882.39 1.27 0,0 9,365.89 1.068 1,0

13,125.1 0.7619 b' I't <-X3I';


Table 2 b. Absorption bands of rninor atrnospheric eonstituents present on line eornpilation [44]

Vo Molecule/ Band ES,' Range No.of (ern-I) isotope (ern-I) Iines

v' VIF

NO 46 0 0 3.217 0.798 5- 98 112 1,875.972 NO 46 1 0 503.4 124.825 1,673-2,072 637 3,723.853 NO 46 2 0 8.301 2.058 3,521-3,901 550

S02626 000 000 239.0 59.255 1- 175 4132 517.75 S02626 010 000 389.9 97.677 433- 617 3326

1,151.7135 S02626 100 000 351.9 87.255 1,048-1,262 5812 1,362.030 S02626 001 000 3,080. 763.673 1,316--1,394 2075 2,499.8701 S02626 101 000 71.37 17.695 2,451-2,528 2075

749.654 N02 646 010 000 370.9 91.968 600- 900 4594 1,616.852 N02 646 001 000 6,113. 1,515.625 1,550--1,657 3276 2,906.073 N02 646 101 000 257.9 63.947 2,833-2,937 1586

949.878 NH3 4111 0100 0000 2,151. 533.403 609-1,266 721

HN03 146 Pure rot 497.1 123.259 10-- 38 66 897. HN03 146 2V9 190.0 47.103 891- 899 1079

1,324.9 HN03 146 V4 2,744. 680.430 1,326-1,335 24 1,709.568 HN03 146 V2 1,590. 394.129 1,719-1,730 1014

3,569.6398 OH61 0 173.1 42.922 3,086-4,011 166

HF 19 0 0 5,704. 1,414.204 41- 589 15 3,961.4429 HF 19 1 0 1,547. 383.671 3,381-4,339 25 7,750.7949 HF 19 2 0 49.62 12.302 7,143-7,993 22

Hel 17 0 0 262.0 64.969 21- 382 19 Hel 15 0 0 806.8 200.042 21- 382 19

2,883.8850 Hel 17 1 0 147.2 36.495 2,486--3,136 33 2,885.9765 Hel 15 1 0 452.8 112.266 2,459-3,138 34 5,663.9276 Hel 17 2 0 3.48 0.863 5,304-5,824 27 5,667.9832 Hel 15 2 0 10.71 2.655 5,272-5,830 29 8,340.9407 Hel 17 3 0 0.02 0.006 8,125-8,448 18 8,346.7771 Hel 15 3 0 0.07 0.018 8,059-8,454 21

HBr 11 0 0 234.6 58.163 17- 338 21 HBr 19 0 0 239.8 59.452 17- 338 21 HBr 11 1 1 65- 129 5 HBr 19 1 1 65- 129 5

2,558.5308 HBr 11 1 0 72.32 17.932 2,195-2,772 36 2,558.0105 HBr 19 1 0 73.92 18.327 2,195-2,773 36

5,026.6005 HBr 11 2 0 0.75 0.187 4,712-5,160 28

5,027.3408 HBr 19 2 0 0.77 0.191 4,713-5,160 28

7,404.1928 HBr 11 3 0 0.02 0.005 7,205-7,494 20

7,405.2610 HBr 19 3 0 0.02 0.005 7,206--7,496 20

9,690.9914 HBr 11 4 0 0.01 0.002 9,506-9,757 18

9,692.3579 HBr 19 4 0 0.01 0.002 9,508-9,759 18

HI 17 0 0 106.7 26.454 13- 286 23

HI 17 1 1 0.002 50-- 137 8

2,229.5817 HI 17 1 0 1.758 0.436 2,118-2,398 26

4,379.2261 HI 17 2 0 0.633 0.157 4,117-4,489 32 6,448.0348 HI17 3 0 0.343 0.085 6,176--6,520 30 8,434.7076 HI 17 4 0 0.059 0.015 8,190--8,487 26

162 H.-J. Bolle

Table 2 b (eontinued)

Vo Moleeule/ Band L"s;a Range NO.of (ern-I) isotope (ern-I) Iines

r' u"

835.480 CIO 76 0 10.02 2.485 803- 865 131 842.554 CIO 56 0 34.87 8.646 796- 879 175

858.9677 OCS 622 1000 0000 109.8 27.220 818- 891 181 2,062.2 OCS 622 0001 0000 7,881. 1,954.116 2,016-2,089 181

2,500 H 2 CO 126 000002 000000 4.55 2,744-2,812 5 2,655 H 2 CO 126 001100 000000 0.59 2,734-2,735 1 2,716.156 H 2 CO 126 001001 000000 138.6 2,700-2,878 105 2,782.457 H 2 CO 126 100000 000000 828.8 205.494 2,723-2,842 424 2,843.326 H 2 CO 126 000010 000000 974.2 241.556 2,704-2,982 595 2,905 H 2 CO 126 010100 000000 38.13 2,734-2,999 28 3,000.066 H 2 CO 126 010001 000000 8.55 2,897-2,957 3

a Surn of the line intensities on eornpilation (first eolurnn in units of 10- 20 ern/rnoleeule, second colurnn in crn- 2/atrn) at 296 K

Table 2c. Ernpirical constants to cornpute the equivelent width for water vapor and carbon dioxide bands after Eqs. (78) respectively (79) after Howard et al. [46J

Gas Band Wavenurnber c k C 0 K Transition equiva-(J.UIl) interval (ern - I) lent width crn - I

H20 0.94 10,100-11,500 38 0.27 -135 230 125 200 1.1 8,300- 9,300 31 0.26 -292 345 180 200 1.38 6,500- 8,000 163 0.30 202 460 198 350 1.87 4,800- 5,900 152 0.30 127 232 144 257 2.7 3,340- 4,400 316 0.32 337 246 150 200 3.2 2,800- 3,340 40.2 0.30 -144 295 151 500 6.3 1,150- 2,020 356 0.30 302 218 157 160

HDO 3.7 2,670- 2,770 0.325 0.37

CO2 1.4 6,650- 7,250 0.058 0.41 80 1.6 6,000- 6,550 0.063 0.38 80 2.0 4,750- 5,200 0.492 0.39 -536 138 114 80 2.7 3,480- 3,800 3.15 0.43 -137 77 68 50 4.3 2,160- 2,500 - 27.5 34 31.5 50 4.8 1,980- 2,160 0.12 0.37 60 5.2 1,870- 1,980 0.024 0.40 30

15.0 550- 800 3.16 0.44 68 55 47 50

dependence. Only purely empirical relations can be established. Fortunately there have been a number of reliable measurements from which the absorption coefficient can be derived. Very prominent is its dependence on the water vapor partial pressure e.


The continuum absorption coefficient can be represented by the following formula [50, 51a]

where e =water vapor partial pressure p = total pressure WH,O = water vapor surface density per unit path length (kg m - Z rn-I)

C(v, T) = self broadening coefficient for water vapor

163

(80)

y = ratio of foreign broadening coefficient for water vapor (predominantly due to nitrogen) to se1f broadening coefficient.

In the 8-14 ~ region the coefficients have the following values:

C(v, T)=C(v,296K)exP{1,800(~- 2;6)}

and with v in cm - 1

C(v, 296 K) =4.18 + 5,578 exp( -7.87.10- 3 v) g-l cmz atm- 1

=4.13 + 5,505 exp( -7.87.10- 3 v) mZ kg- 1 (MPa)-1

an alternative formula is

C(v, 296 K)=4.13 + 7,729 exp( -8.30.10- 3 v) mZ kg-1(MPa)-1

(81)

(82)

(83)

which extends into the 30 ~ region with slightly less accuracy in the 8-14 ~ region then Eq. (82). There is some new evidence that the values of C(v, 296 K) as given above may be too high by up to 30%.

The value for y is neither very certain nor extremy important. Values of 0.001--0.002 are generally accepted, though maximum values of 0.005 are also reported.

In the 3.5-4.2 I-lm region the water vapor continuum can be expressed by

C(v, T)=C(v, 296 K) eXP[4.56(2~ -1)] (84)

with the values for C(v, 296 K) following in Table 3. In recent publications Carlson [51 b, c] reported that the quasi-continuum absorption of water

vapor is proportional to the ion content of moist air. The ions can be formed by dissociation of water molecule clusters (HZÜ)m+n+l--->H+(H2Ü)m+üH-(H2Ü)n or HX(H20)m--->H+(HzO)m+X(X- = negative ion or electron).

Table 3. Continuum absorption coefficient for 296 K in the 4 ~ region

2,400 2,440 2,500 2,550 2,600 2,650 2,700 2,800 2,900 3,000

C(v, 296 K) in cmz g-l atm- 1

0.1877 0.155 0.117 0.099 0.090 0.101 0.120 0.168 0.237 0.328

0.185 0.152 0.115 0.098 0.089 0.100 0.118 0.166 0.234 0.324

164 H.-J. Bolle

Collision Induced N 2 Absorption

Nitrogen normally has no vibrational absorption because it does not have a permanent electric dipole. Due to a collision induced dipole moment there exists, however, an absorption continuum between 2,075 and 2,700 cm- 1 (3.7-4.81lffi). Due to the nature of its generation the absorption coefficient is proportional to the square ofthe press ure [52].

Electronic Spectra

Electronic transitions are generally only excited ifthe energy exceeds the equivalent of about IIlffi wavelength (1.99· 10- 19 J or 1.22 eV). Molecular magnetic dipole transitions occur already at 1.2683 Ilffi and 1.0674 Ilffi as weIl as at 0.7620).1l11 and 0.6317 Ilffi (red "atmospheric" bands of oxygen, see Table 4). Of great importance are the electronic bands of ozone since they absorb the biologically dangerous uv radiation in the Huggins bands, 310-360 Ilffi. Other prominent electronic bands are the Schumann-Runge bands of O 2 (175-203 nm) [53, 54].

Dissociative Absorptions

Molecular oxygen starts to absorb in the ultraviolet very weakly at about 250 nm. The first dissociation limit is reached at 242.3 nm where two 0 ep) atoms are produced, and where the weak Herzberg continuum (203-240 nm) starts. At 202.6 nm the strong Schuman-Runge bands begin and become dissociative at 175 nm where the transition into the Schumann-Runge continuum (125-175 nm) occurs. Absorption of these higher frequencies result in the production of one oxygen atom in the ground state [Oep)J and one in an excited state [Oe D)J. Superimposed to the Schumann-Runge continuum are some bands which indicate dissociation products in even higher excitation levels.

For nitrogen the dissociation threshold is reached at 127 nm near the short wave end of the Lyman-Birge-Hopfield bands which start at 145 nm and extend down to 112 nm. No strong nitrogen dissociation continuum has been observed, nor nitrogen electronic bands above 145 nm.

Ozone starts to dissociate already at wavelengths near 1.11lffi, where the Chappius continuum (0.43-1.1 ).1l11) has its longwave end.

Some of the more important dissociative reactions and wavelength limits are listed in Table 4 (Compare also Fig. 10).

Photo-I onization

Photo-ionization potentials are generally higher than the dissociation limits. The wavelengths range around 100 nm for the most important species, more specifically: from 85.3 nm for N(4S) to 134.0 nm for NO(X2 II). The ionizing radiation is therefore partly screened by the molecular oxygen absorption, except for narrow "windows" for which that one at 121.57 nm is the most prominent one. This wavelength coincides with the strong Lyman-ct emission of hydrogen in the solar atmosphere. Since also CO2 has only a small absorption in this region (118-123 nm) the Lyman-ct radiation penetrates deep into the mesosphere and ionizes NO in the so called D-Layer down to about 70 km (compare Fig. 10).


Table 4. Seleeted photo dissoeiation proeesses of atmospherie species. (After [39, 53-56]). Symbols:

a)Atoms Orbital angular momentum= 1 Quantum States I=O:S, 1=1:P, 1=2:0 Spin multiplicity=(2S+ 1)=left side upper index

b) Moleeules Quantum number of eomponent of eleetronic orbital angular momentum veetor along axis = A Quantum states A=O:1', A= 1:II, A=2:L1 Parity: g (symmetrie) or u (antisymmetrie) with respeet to inversions. Symmetry with respeet to mirroring: + or -

165

X stands for the ground state if several electronie states are known, others are distinguished by A, B, ...

Atmospherie moleeule

NO

CFCh

CF2Ch

CF3C1

Dissociation produets

20ep) OeO)+Oep)

Oep) + 02(X31';) Oep)+02eL1 g)

Oep) + 02(b11' J 0(lD)+02(L1J OeO)+02(X31'.)

CO(X11')+Oep) CO(X11')+OeS) CO(X11')+OeO) CO(A 31') + Oep)

HeS) + OH(X2ll) H+OH(A21'+) H2e1')+O

N2e1':)+Oep) N2e1':)+OeO) N2e1':)+OeS) NeO) + NOe ll)

OH+N02 HN02 +0ep) HeS)+N0 3

HN02+OeO)

NeO)+Oep) N(4S)+Oep)

OH+NO H+N02

H+HCO H2+CO

O+OH

CH3+H CH2+H2

CFCh+CI

CF2CI+CI

CF3+CI

Oissociation wavelength treshold or wavelength region in nm

242 175

1,180>,1.>450 610>,1.>313 462 313 410

108 128 167 108

200 136 147

740 . 340 210 170

598 401 290 245

190 135

475 366

340>,1.>250 372

-300

280 258

214

213

200

166 H.-J. Bolle

Emission Under Thermodynamic Equilibrium Conditions

According to Kirchhoff's Law the emission from a body equals the radiant exitance of a black body times the absorption coefficient of the (non black) material:

M=aMB(T). (85)

According to Planck the spectral black body exitance, its non-polarized emission from an unit area into the hemisphere, is a function of the temperature only, and is represented in dependence of the wavelength by the expression.

and in dependence of the wavenumber by

MB,v(T)=C1V3 (ec2V/T_1)-1

with Cl =211: hc2 = 3.7418· 1O- 16 W m -2 first radiation constant C2 =hc/k= 1.4388 .1O- 2m K second radiation constant h = 6.6262 . 10 - 34 J s Planck constant k = 1.3807 . 10-2 3 J K - 1 Boltzman constant c = 2.9979 . 108 m s - 1 speed of light in vacuum.

(86)

(87)

Figure 8 gives the exitance for the wavelength range 1-100 J.Ull and the temperature range 50-6,000 K. Accordingly the spectral radiance LB,,, of a black body normal to the radiating unit area, for which the symbol B,,(T) will be used, is

(88)

Its unit is W m - 2 sr - 1 (cm - 1) - 1 respectively W m - 2 sr - 1 J.Ull- 1 if the quantity B ;.(T) is used.

In the atmosphere the basic proposition for black body emission, the thermodynamic equilibrium, is, strictly speaking, not completely achieved. Photons emitted at one altitude of the atmosphere will interact with molecules at another altitude where the population of the different involved rotational energy levels may be slightly different, because of a difference in the temperature due to the vertical atmospheric temperature gradient. For the atmosphere the expression "Ioeal thermodynamie equilibrium" (LTE) is commonly used if one refers to a situation where deviations from the thermodynamic equilibrium are negligible. At lower pressures the transfer of absorbed radiant energy into translational energy due to collisions takes longer because the mean free path length of a molecule increases. Careful studies of this problem have resulted in the conclusion that deviations from LTE can be neglected in most cases up to altitudes of '" 70 km but become essential for the thermosphere [39].

For a horizontally infinite column of atmospheric gases the exitance can be derived in the following way. An infinitesimal path ds' in the column emits proportional to the radiating mass (1(s) ds' [(1(s') is the mass area density measured in kg m - 2], and a so far unknown source function Xv<, T). The spectral radiance emerging from ds' around s' can be written (Fig. 9):

d4(s')=(1(s')X,,[T(s')J ds' . (89)


Fig. 8. Black body exitance. Emission of black body into the hemisphere for temperatures between 50 and 6,000 K and 1-100 nm wavelength

167

106 r----------,

Wavelength (pm)

168

~ ~ ~ I I I

s-• .-------------~---------------------I-- dS'--1

i~

q (5') dUs')

~~ ~ = X [T(s')] ~(s') ds'

Fig.9. Derivation of column radiance, Eq. (92)

1--L=_-X[T],...;I\\,Il[;_(S',s .... O) ds'

= -X[T]e(O-I)

H.-J. Bolle

Radiance emerging from elementary layer ds

Radiance from elementary layer after passing layer 5' ...... 50 with transmittance 'T' (5', 50)

Radiance emerging at So from infinite column of constant temperature and radiative properties

At an observer posted at So arrives only a fraction of the radiation emitted at s' because of the absorption between s' and So. According to Beer's law this fraction is given by the transmission function:

r(v; s', so)=exp {-i am(v; s) gIs) dS}. (90)

The integral in the r.h.s. exponential function is the (spectral) optical path length u,,: S·

u(v; s', so) = J am(v; s)g(s) ds. (91) So

The absorption coefficient a m is here defined per mass area density (mass absorption coefficient). From an air column between U(SI) and u(so=O) the following radiance is emerging (for the

moment the spectral index vwiIl be dropped):

L(SI, so) = - Si r(s', so) dL(s') o s, -le(s)am(s)ds

= - J eIs') X[T(s')] e 0 ds' o

= _ l' X [T(s')] (dr(S', so») ds' o am(s') ds'

,(s, • sul X [T(s)] = - J --(-)-dr.

1 am S

(92)

If we now assume that the air column is enclosed in a black body of constant temperature T of semiinfinite extension, the expression (92) must equal the radiance of a black body at temperature T for r(sl, so)->O:

L(Sl-> 00, So =0) = LB(T) = B(T) . (93)


For T=constant and am(s)=constant Eq. (92) can be integrated between so=O and s=oo with ,(00,0)=0 and ,(0,0)= 1 and it results

(94)

or

(95)

With Eq. (95) the general result has been obtained that in agasous medium the source function for the radiance L = Mire is equal to the Planck function times the absorption coefficient as long as L TE can be presumed.

Non-Thennal Emissions in the Upper Atmosphere

At reduced pressures with increased free path lengths the quenching of excited states due to collisions is reduced. The times between collisions (10- 10 s at the surface, 10- 5 sand 10- 4 s in 75 km resp. 90 km altitude) may become longer than the life time of excited molecules and atoms. In this case the molecule or atom will emit one or more photons to get back into its ground level. In the course of this process even transitions may occur which under other circumstances are forbidden by quantum theory.

Often the excited species are dissociation products. Table 5 compiles the most prominent emissions which are observed as airglow [57-60]. Other emissions which are correlated to the intensity of the atomic oxygen green line (557.7 nm) are from molecular oxygen and include the Herzberg bands in the near ultraviolett and in the blue part of the spectrum, the Chamberlain bands in the blue, a continuum in the green and an emission at 864.5 nm which belongs to the "atmospheric" band system (Kaplan-Meinel bands). The total emission of these bands is about 3 kR. Rayleigh (R) is an often used unit for the rate of photon emission per unit area in a column (s -1 m - 2). For conversion into SI Units see Table 5. The conversion becomes difficult in case of structured broad band features. The emission orginates around 100 km altitude.

The "normal" airglow emission is strongly enhanced at high latitudes during solar activity which generates the aurora. Other emissions partly from highly excited atoms are at that time observed in addition. These emissions are excited by high energy electrons and protons which are emitted by the sun during eruptions or captured in the Van Allan belts of the earth from where they are discharged into the atmosphere.

Atmospheric Radiation Field

Solar Radiation

Extraterrestrial Solar Spectrum

The problem to measure the extraterrestrial solar spectrum and its possible variations is not yet solved satisfactorily. The corrections which have to be applied to ground based measurements due to the atmospheric extinction in the ultraviolett

->

0

Tab

le 5

. P

rom

inen

t no

n-th

erm

al a

tmos

pher

ic e

mis

sion

s an

d th

e 0

1 f

ar i

nfra

red

emis

sion

(1 R

~ 1

.58)

" -1

. 1

O-'

6W

m-2

sr-

" A

in m

)

Spec

ie

Tra

nsit

ion

Wav

elen

gth

Hei

ght

rang

e In

tens

ity

(kR

) A

ppro

x. r

adia

nce

Pro

duci

ng

lJ1ll

km

Wm

-2 s

r-1

Day

N

ight

R

eact

ion

Wav

elen

gth

(nm

)

N

2p2D

--->2

p4S

0.51

99

150-

280

0.05

N

2 +e-

--->N

+N

eD

)+e

80-1

00

N2

+hv

--->

N+

NeD

) N

O+

+ e

--->

NeD

) + 0

u.

a.

0 'S

--->

'O

0.55

77

90-2

20

3 0.

25

0.7

-8.5

.10

-7

0+

0+

O--

->0

2+

0('

S)

<2

42

<

313

313-

350

0 '0

2---

>3P

2 0.

6300

10

0-40

0 2-

20

<

0.5

1.3

-50

.10

-7

O2

+hv

--->

O('

D)+

Oep

) 'D

2--->

3P1

0.63

64

03

+ h

v---

>O('D

) + O

2(' d

g)

O2

b1.r

: --->

X3 .r

: 0.

7619

(0-0

) <

50

-100

30

0 6

O.l2

-{)

. 10-

5 0

3 +hV

--->

02(b

1 .r:)

+ O

ep

) 0.

8645

(0-

1)

NaO

+ O

--->

Nae

p) +

O2

Na

3p2p

~,,---

>3s2

St

0.58

90

70-1

10

0.H

l.3

N

aH+

O--

->N

aep

)+O

H

0.58

96

NaH

+H

--->

Nae

P)+

H2

O2

a'd

g---

>X3 .r

;-1.

270

(0-0

) 40

-85

2.1

04

80

0.0

96

-25

.10

-4

03

+ h

v---

>Oep

) + 0

2(a

'dg

)

1.58

(0-1

)

OH

V

ibr.

ban

ds

0.38

2-4.

500

55-9

5 3

.10

3 10

3 1

-3.1

0-4

H

+0

3---

>O

H+

02

v:S:

9 H

02+

O--

->O

H+

02

0 3P

1---

>3P 2

63

.09

} 10

0->

200"

) T

herm

ally

exci

ted

0 3P

O---

>3P 1

14

7.06

0

p:: , :-<

" P

rodu

ces

heat

ing

betw

een

80-1

00 k

m

tc e.. "


and infrared can not be determined to the desired accuracy, and measurements from high flying aircraft and space platforms outside the atmosphere suff er from inadequate control of their absolute calibration.

Two major programs have been carried out in the past which started from different points ofview. Labs and Neckel [61, 62J made carefully calibrated high spatial resolution measurements of the solar radiance at the center of the sun from the ground (Jungfraujoch, 3,500 m) and corrected the measurements for solar limb darkening and atmospheric attenuation. Thekaekara [63, 64J evaluated measurements of the spectral solar irradiance made from high altitude aircrafts, balloons and space platforms. In these experiments attempts have been made to measure the radiation flux from the whole sun at 1 Astronomical Unit, the mean radius of the earth's orbit.

The state ofthe art of solar spectral irradiance measurements was reviewed at a workshop and a symposium held in Washington [65, 220J, by White [66aJ, and more recently by London and Fröhlich [66b]. A proposed standard curve of solar spectral irradiance as deduced by Thekaekara is referred to in different publications [67, 68, 69J, where the interested reader will find the data in tabulated form. Two words of caution in using these data are necessary:

a) The discrepancy between the proposed spectral distributions and that one of Labs and Neckel in their narrower spectral interval (401~57 nm) is not yet resolved, and

b) the proposed spectral distributions are normalized to a solar constant of 1,353 Wm- 2 which is probably 15-17Wm- 2 too small. The Fig. 10 presents the short wave part of the spectrum in graphical form

together with the penetration depth into the atmosphere. For the solar spectral irradiance at the top of the earth atmosphere in

1 Astronomical Unit (A. u.) = 1.496 . 1011 m distance from the sun the symbol So, A is introduced. 1 A. U. is the mean radius of the earth orbit around the sun. The total flux

00

So= !So,AdA,=1370Wm- 2 (96)

is called Solar Constant and will be discussed later. The electromagnetic wave energy reaching the earth orbit at mean distance

from the sun per unit area is traditionally called the Solar Constant (So) though it is not a constant in a physical sense. The solar electromagnetic radiation is not produced by a simple black body emission but is generated in different layers ofthe solar atmosphere which has a wide temperature range. Far infrared radiation around 100 J.UIl wavelength is emitted by the relatively cold upper photosphere with a temperature range of 4,150-4,450 K. The near infrared emission indicates temperatures between 6,000 and about 6,700 K. The maximum, around 1.6 J.UI1, originates in the deepest observable layers of the solar atmosphere. The visible part of the spectrum can be approximated by a 6,000 K black body radiation which corresponds to temperatures in the middle photosphere. This part of the spectrum is, however, modified by the absorption of gases in the higher and colder layers ofthe photosphere which causes the dark Fraunhofer lines. In the optically

172

.. u c:

'" -0

~

E .:.:;

.J::.

C. .. -0

c: .S!

'" Gi c: ..

Il..

\0-3

120

100

80

60

40

20

01

H.-I. Bolle

12r----------------,

Wavelength l,u m 1

Ratio Labs & Neckei : NASA

Exlraterestrial solar irradiance

"NASA

\ ~. I

~\ !I A t\

\ l' J • I. \

'/ ' 1 , , " " ~ ,

~/

Penetration depth "lA \l

Wave length (nml

Fig. 10. Solar irradiance at the top of the atmosphere at I A.U. distance between sun and earth after NASA [67] and NOAA [69]. Lower part: penetration depth for which the radiance is attenuated to Ije. (After Friedman [70])


thin cromosphere temperatures rise again to several ten - to hundred thousand degrees which produce lines from ionized atoms in the shortwave ultraviolett [70]. Finally in the corona the temperatures are larger than 106K.

Attenuation of Solar Radiation in the Atmosphere

The spectral Rayleigh scattering coefficient O'R(A) was derived earlier [Eq. (26)] and defines the attenuation per unit length. In Table 6 some values are presented for the conditions at the bottom of amid latitude standard atmosphere (T=288.15 K, p= 1.01325.105 Pa) according to [18]. In order to determine the optical path through the whole atmosphere one has, in principle, to integrate Eq. (21) from the top to the bottom of the atmosphere along the line of sight.

In order to derive a formula for the atmospheric transmission let us assume that the sun is located at a zenith angle C. Until it reaches the ground the solar radiation has to pass a larger air mass then for vertical incidence. A relative airmass M (C) can be defined which gives the ratio of the airmass traversed under the zenith angle C to vertical incidence. Because of the curvature of the earth and refraction the relative airmass is a complicated function ofthe zenith angle [71, 72].

Table 6. Rayleigh scattering coefficient, optical depth and transmittance as weil as the turbidity factor for amid latitude standard atmosphere [18]

Wavelength Rayleigh T(A) Rayleigh Transmission due to in J.lID scattering optical depth Rayleigh scattering only

coefficient "R(A) at STP in m- 1 (0=0° (0=75°

0.27 2.282.10- 4 (38) 1.9314 0.1449 0.00057 0.28 1.948.10- 4 (23) 1.6481 0.1924 0.00172 0.30 1.446.10- 4 (4.06) 1.2237 0.2941 0.0209 0.32 1.089.10- 4 (1.67) 0.9290 0.3949 0.0276 0.34 8.494.10- 5 (1.46) 0.7188 0.4873 0.0622 0.36 6.680.10- 5 (1.54) 0.5653 0.5682 0.1126 0.38 5.327.10- 5 1.65 0.4508 0.6371 0.1752 0.40 4.303.10- 5 1.70 0.3641 0.6948 0.2449 0.45 2.641.10- 5 2.03 0.2238 0.7995 0.4212 0.50 1.726.10- 5 2.55 0.1452 0.8648 0.5706 0.55 1.162.10- 5 (3.36) 0.0984 0.9063 0.6837 0.60 8.157.10- 6 (4.42) 0.0690 0.9333 0.7660 0.65 5.893.10- 6 (5.05) 0.0499 0.9513 0.8246 0.70 4.364.10- 6 (5.88) 0.0369 0.9638 0.8671 0.80 2.545.10- 6 (8.70) 0.0215 0.9787 0.9203 0.90 1.583.10-6 12.48 0.0134 0.9867 0.9495 1.06 8.458.10- 7 21.0 0.0072 0.9928 0.9726 1.26 4.076.10- 7 41.5 0.0034 0.9966 0.9869 1.67 1.327.10- 7 114 0.0011 0.9989 0.9958 2.17 4.586.10- 8 273 0.0004 0.9996 0.9985 3.50 6.830.10- 9 890 0.0001 0.9999 0.9996 4.00 4.002.10- 9

174 H.-J. Bolle

Table 7. Relative air (M) and water vapor (Mw) masses in dependence ofthe zenith angle'

75 80 81 82 83 84 85 86 87 88 89 90° 3.81 5.59 6.16 6.86 7.73 8.85 10.32 12.33 15.18 19.46 26.31 36.26 3.85 5.71 6.33 7.09 8.07 9.35 11.1 13.7 17.6 24.5 38.6 75.1

However, for (~80° it can be approximated by

(97)

Por r:?75° the correct values are given in Table 7. In this table also the relative water vapor mass Mw(() has been included.

Because of the different height distribution of water vapor the sec ~ condition holds much closer.

The transmission of the atmosphere can now be described by the general formula

-IiM(1;) -IiRTM(() 't'=e =e (98)

where () is the total optical depth of the atmosphere for vertical incidence (sun in zenith), and ()R the respective value for Rayleigh scattering. T is the "turbidity factor" introduced by Linke [73]. In the formulation of Eq. (98) it is assumed that the transmittance can be computed from the molecular scattering approximation if the optical path is prolonged by a factor T. T is essentially the number of Rayleigh atmospheres which have to be put on top of each other in order to produce the same extinction as the turbid atmosphere.

Equation (98) can first be considered for narrow, nearly monochromatic spectral bands. In this case the spectral optical depth and the spectral turbidity factor T(l) have to be used in the computations. Typical values of T(l) for a clear standard atmosphere are also given in Table 6. The values in parenthesis are for spectral regions in which atmospheric gaseous absorption is present (mainly ozone). It can be seen that the turbidity factor increases rapidly towards longer wavelengths because of the much different wavelength behavior of molecular and aerosol scattering. The molecular scattering decreases proportional l -4.09 and the aerosol scattering proportional", l -1.5. The various effects which contribute to the attenuation of the solar radiation are schematically presented in Pig. 11. Since the molecular scattering can easily be computed from the known Rayleigh scattering coefficient Eq. (26), a measurement of the spectral transmittance in combination with Eq. (98) provides a measure for T, the turbidity of the atmosphere.

Measurements of the spectral atmospheric transmittance are of high practical value for pollution monitoring. Prom these measurements not only can the total aerosolload of the atmosphere be deduced but also information about its size distribution be derived. The scattering efficiency at a specified wavelength of aerosols depends on its size distribution. Therefore the spectral dependence


N

'E 3 .:x

Q) <.l C

"'

2

~ , ~ ti Q)

Co (/)

Sea t tered by motecules

A bsor bed by Olone

Scattered by aerosot partlcles

Extraterrestflal i rradlanee

Absorbed by water vapof

by ozone

Wavelengl h (pm)

Fig. 11. Attenuation processes for solar radiance in a cloudfree atmosphere (After Quenzel [74])

175

of the scattering coefficient contains information on this size distribution which can be inferred by the application of mathematical inversion methods [75, 76].

By integration over the whole solar spectrum the spectrally averaged turbidity factor can be deduced. It has, however, to be taken into account that ozone and water vapor also ab so rb and thus contribute to the extinction (turbidity va lues contaminated by water vapor or ozone absorption are put into parenthesis in Table 6). Furthermore the Rayleigh optical depth gets dependent on the zenith angle of the sun because with increasing zenith angle the transmitted solar spectrum is shifted towards longer wavelength according to the wavelength dependence of the scattering coefficient (see next section).

To arrive at reliable aerosol turbidity values from total solar radiation measurements is not a straight foreward procedure and requires experience (see e.g. Robinson [77]). Some guiding va lues after Schulze [78] are:

High Mountain area Rural Lowland Large cities Industrial area, highly poIlu ted

T=1.9 T=2.75 T=3.75 T=5.0.

More detailed data are given by Linke [73]. In a cloudfree atmosphere with low turbidity the daily average transmission

varies over the year between 0.84 and 0.93. The average relative irradiance is 0.77±0.01 at the equator with a mean solar zenith angle of (=52.3°, and varies between zero (polar night) and 0.69 at the poles.

176 H.-J. Bolle

Deduction oJ the F ormula Jor the Transmittance oJ a Scattering Atmosphere

We consider the sun to be at zenith angle ,. Instead of the geometrical path s through the atmosphere the vertical height z can be introduced

and with the relative air man M(,) the geometrical path becomes

ds=M(O dz. (99)

Using Eq. (99) the optical path UR (A, ,; zo, Zoo) through the whole atmosphere can be expressed for molecular scattering by

00

UR(A, ,; O,oo)=! O"R(A, z) M(O dz, (100)

where O"R(A, z) is the Rayleigh scattering coefficient for the pressure and temperature conditions at z. The Rayleigh scattering coefficient is defined by Eq. (26) for N scattering centers per unit volume.

The number of molecules is defined by the equation of state for an ideal gas

N=p/kT=N AfI/RT=7.243 .1022 f ~a (101)

N A=6.022 .1023 mol- 1 is the number of molecules per mol (Avogadro's number), R=8.314J mol- 1

K - 1 the molar gas constant and T the absolute temperature. The formula for the Rayleigh scattering coefficient is therefore in terms of pressure and

temperature given by

Ainm. For p = 105 Pa, T=273.15 K and (n -1) computed after Eq. (31) it results for A= 550 nm

327[3(2.911 .1O- 4 f ·273.15 O"R(550 nm)= 3.7.243.1022 .105(0.55.10 6)4 (103)

=1.16·1O- 5m- 1 .

Let us assume vertical incidence (p.= 1) first In this case the symbol Ö is introduced for the vertical optical path

which is called the optical depth of the atmosphere. More generally

Ö(A; Z)=U(A, '=0; z, 00)

is the optical depth of the atmosphere at the altitude z. The optical depth of the entire atmosphere due to Rayleigh scattering,

00

ÖR(A, 0) = - J O"R(A, z) dz, o

(104)

(105)

(106)

can in good approximation (about 1.5% error) be expressed by the value O"R(A, 0) at STP conditions and the equivalent height H of the atmosphere which would· be its vertical extent if the whole atmosphere is compressed to 0.1013 MPa pressure at 273 K:

(107)

H is the scale height of the atmosphere and can be computed from the hydrostatic equation

gedz= -dp and the state equation of~n ideal gas, p=e ~ , as . H=RTo ~8000m

M.g- , (108)

(R = 8.314 J mol- 1 K - \ molar gas constant; To = 273.15 K; M. = 0.028964 kg mol- \ molar mass ofair up to about 90 km; 9 = 9.8062 m s -1, acceleration due to gravity at 45° latitude).


With O"R(550nm,0)=1.162.1O- s m- 1 from Table 6 and the value of H from Eq. (108) the optical depth of the atmosphere due to Rayleigh scattering at 550 nm is according to Eq. (107)

oR(550 nm, 0) ~0.093 (109)

which is equivalent to a transmittance of

'R(550 nm, M(O= 1; 0, <Xl)=e- 0 .093 =0.91. (110)

Thus 9% of the energy is scattered out of the direct beam. The transmittance for an inclined beam is according to Eq. (100)

(111)

Values for the Rayleigh optical depth and transmittances for M(O=1 and M(O=3.82 presented in Table 6 have been computed for amid latitude standard atmosphere (To = 288.15 K) under consideration of the true temperature profile [18].

Now the spectral irradiance of the sun at the earth surface under normal incidence can be computed by application of Eq. (111):

S ( -O)-S -öRP.,O)M(() ,;,z- - o,J,e . (112)

The total solar irradiance for a zenith angle *0 at the horizontal earth surface, EQ , can be obtained from Eq. (112) by multiplication of the unit area by cos (0 = fl and integration over all wavelengths.

Ws - Ö.(,;, O)M(C) . EQ(fl)=fl So,;,e d2=flSo'R' (113)

o

As indicated in Eq. (113) it is desirable to express the solar irradiance at the surface by the solar constant defined by Eq. (96) and an integrated transmittance 'R(O)

'_ -ÖRM(O_ Ws So.;, -ÖR(A,O)M(Od' 'R-e - --e /c,

o So (114)

Due to this definition the spectrally integrated optical depth 0R becomes dependent on the relative air mass M((). Numerical integration of equation (114) leads to the following representation of 0R:

(115)

Values obtained for pro are given in Table 8. With increasing zenith angle the spectrally averaged and vertically measured optical Rayleigh

depth of the atmosphere decreases because in the spectral distribution of the transmitted radiation the longer wavelengths where more energy is contained (compare Fig. 10) are favoured due to the 2 -4.09 dependence of the Rayleigh scattering coefficient. The spectrally integrated optical depth for vertical incidence [M(O= I] has the numerical value oR(I) = 0.0992 which results in a transmittance ofr = 0.906.

Aerosol extinction can be included in the computation by generalizing e.g. Eqs. (111) and (106) to

,()" (; z, <Xl) = exp {-1 [O"R(A, z')M(C z') + O"M(2, Z')MM(C z')] dzJ (116)

Because of the different vertical distributions the relative air mass for aerosols is not exactly the same as for molecules, Also M(O is not constant with height because of the curvature of the earth (see Fig. 12). Again for angles « 75°, M(C Z)=MM((, z)= sec ( is a sufficiently good approximation. Then

Table 8. Values of the reduction factor P(() defined by Eq, (115) in dependence of the relative air mass M(O

M(() pro

0.5 1.058

1.0 1.0

1.5 0.949

2 0.900

3 0.826

4 0.768

6 0,675

8 0.608

10 0.555

178 H.-J. Bolle

I --I ~-----

.~

sin~' = _r_ sint f+Z

ds dz dz = cos ~' = -:ij/=1_=j(~fE::='T.)2;=s=i=n ~~2

f E+ Z

Fig. 12. Optical path in a spherical atmosphere

it may be written (under omission of the spectral index):

O"R+O"M= O"R(l + ::)= TO"R· (117)

T is the turbidity factor introduced by Linke [73]. It can be computed spectrally from Eqs. (26) and (65) if the aerosol content and its size distribution are known. It can also directly be deduced from measured transmittances since the Rayleigh scattering coefficient can easily be computed. From Eq. (116) it would follow with MM=M: T= -In r/"RM(O.

However, in this case the turbidity factor would include the effects of absorbing gases. More generally Eq. (116) has to be written:

(118)

where the "(z) denote the spectrally integrated optical depths of the wh oie atmosphere down to a level z. Only if the ozone and water vapor absorptions are known (there is also a small absorption of O2 involved), it is possible to determine the turbidity factor from

(119)


Scattered Solar Radiation

Single Rayleigh scattering, according to it symmetrical phase function, scatters the same amount of solar radiation back to space as in the direction of the earth surface. If So is the solar constant and Edoo)=SoCOSC0 the irradiance at the top of the atmosphere, then the fraction

(120)

arrives at the horizontal earth surface. H is the height of the homogenous atmosphere defined by Eq. (108). The fractional total scattered radiation in a nonabsorbing atmosphere over a black earth surface would then be

1-e -TuRHM((oJ.

A fraction q is scattered toward the earth, generating an irradiance of diffuse radiation of

(121)

In a cloudfree atmosphere according to Berlage [79] 50% ofthe scattered radiation is directed towards the earth surface, 50% are lost to space. If 't'[M(C0)J is the transmission of the atmosphere at airmass M(Cd, then the diffuse scattered radiation is

(122)

For a Rayleigh atmosphere q=0.5 is still a good approximation (",1%) even if multiple scattering is accounted for. In the case of a turbid atmosphere one would expect that because of the stronger forward scattering the factor q will be larger. However, due to multiple scattering this effect is largely compensated.

If one accounts also for the absorption by aerosol, water vapor and ozone in the atmosphere by introduction of the extinction factor aHM(Cd the formula

(123)

represents with reasonable accuracy the irradiance by diffuse clear sky radiation at the horizontal surface of the earth [78]. Different experimental investigations [80,81] lead to the result, that the absorption contributes 10--20% to the total extinction byaerosols. The absorbed radiant energy heats the atmosphere.

The differences in the scattering properties of different air masses can be seen in Fig. 13. Here the zenith radiance is plotted in dependence of the wavelength normalized at 780 nm for the city of Munich at solar zenith angles of 37.3° and 59.0° and for a coastal station in the Mediterranean region at a solar zenith angle of about 43° which is in-between the two values of Munich. It can clearly be seen that the slope of the spectrum for the Mediterranean station is steeper but strong variations occur also in the urban atmosphere. In addition a measurement from a mountain station (Jungfraujoch, 3,570 m) is added which has an even steeper decrease towards longer wavelengths but still not the slope for Rayleigh scattering as indicated by the curve R is reached. In the poIlu ted urban atmosphere there is a tendency towards a more shallow slope in the spectral distribution especially at shorter wavelengths, which is a consequence of the larger aerosol concentration.

180. H.-J. Bolle

OE 10.

=<-N 5 'E 3 3 ... 2 u c <:! -0

f3 0 C, 0::

I 4.0. 10-12

0.5 - 11 5.0. 10-'3 m = 15

0.3

0.2

0. 1 0.5 0.6

Wavelength

Fig. 13. Speetral zenith radianee of seattered solar radiation A: Absolute radianee observed in Munieh, 25. July 1962, (0 = 38°. B:Radianee observed in Munieh, 2. Oetober 1962, (0 = 59°, norrnalized to speetrum A at 780. nrn. C: Radianee observed in S. Agata sui due golfi, <p=40.°37'N, A= 14°11'E, 19 August 1961, normalized to speetrurn A, at 780. nrn, D: Absolute radianee observed at Jungfraujoeh (3570. rn), February 1961. R: Theoretieal Rayleigh seattering eurve for Munieh. H, IH: Radianee eornputed for Mie seattering. (After Bolle et al. [82])

Strong foreward scattering is reflected in high radiance va lues around the sun. The minimum occurs at about 90° off the direction of the sun. Towards the horizon the radiance increases because of the increasing air mass. This has often been measured [82J and can also be computed by application of a rigorous radiation transfer theory as we shall see later (compare Fig. 23).

Atmospheric Longwave Radiation

The radiation which is emitted by the atmosphere and the earth surface due to the temperature structure in the atmosphere-surface system is governed by the Planck function, Eqs. (86) and (87), and the distribution of emitting gases and aerosols in the atnlosphere. For the temperature range wh ich exist in the layers emitting at LTE (180--330 K) the emission starts to reach measurable values at about 2.5 11m and extends through the whole infrared to the mm region. The maximum spectral radiance values per unit wavelength interval occur around 10 11m wavelength (around 600 cm - 1 if plotted per unit wavenumber interval).

The transition between scattered solar radiation and infrared radiation emitted by the atmosphere can be recognized in Fig. 14. In its lower part at the left side the zenith radiance of scattered solar radiation of the clear sky is reproduced


o 0-

c o

~ 50 E U1 C o ,::

Wave[ength (}Im)

0.6 0.8 1.0 1.5 2 3 4

OL---------------~~--~----~~----~~~--~--L-~

(lJ u C o

-cl

10

fi. 0.01

0.6 0.8 1.0 1.5 2.0 3 4 5 6 7 8 9 Wave[englh (}Im)

181

10

.0.1

0.01

Fig. 14. Atmospheric spectrum 0.5-9 /lm. Upper part: Transmission through a horizontal atmospheric path of 5,500 m with 1.37 g cm -2 water vapor (dashed line 300 m path with 0.11 g cm -2 H 20) (After Yates and Tylor [83]). Lower part: Zenith spectral radiance of scattered solar radiation and emitted atmospheric radiation. (After Bolle [84])

with theslopealready known from Fig. 13. Near 2.5 ~ the emission of atmospheric gases starts to overlap the scattered radiance and increases towards longer wavelengths. In strong vibration-rotation bands like the COz band at 4.3 ~ and the water vapor band at 6.3 ~ (compare Table 2a) the emission reaches the Planckian Function (dashed curve) for the air temperature near the surface (285 K). The structure in the "windows" at 3.8 and 4.8 ~ as observed under higher spectral resolution can be seen in the displaced inserts as measured at other occasions [85, 86]. In the upper part of Fig. 14 the atmospheric transmission measured by Yates and Taylor [83] is plotted for comparison. The bands denoted by letters in the scattered solar spectrum are identified as folIows:

A B C

Oz (0-0) 0.7621 ~ Oz (0-1) 0.6884 ~ H~ 0.65628 ~

D NaI 0.58899 ~ 0.58959 ~.

HzO bauds compare Tables 2c and 10

182

, .... '"

'"i'E

3:

'" u c: o "0 o a:

H .-J. Bolle

( ,u m)

25 0.10 .-------;----~-__i:....---.:r_--_T_--__i---=r=_____l

a05

Antarc ti ca 74.6°S, 44.4 oE

o - ...=:-- - - ----

0.15

0.\0

0.05

0.15

0.10

a05

12.00 GMT 14.8° N, 4.7 °W

Stratosphere warmer

thon surface\ __

------ --220K--- --- __ 200K ~ ---------

surfoce tempo

............ -3iOK--- ....... ...... "-

, ...... ------looK ---

Wavenumber

---

"-"'\

--------

400

Fig. 15. The terrestrial emission spectrum as measured from space. (After Hane! et al. [87])


The variability ofthe atmospheric emission is beside ofthe temperature primare!y produced by the variable water vapor mass. The contributions by CO2 and some other minor constituents such as N20 and CH4 (compare Table 2a) are much more constant; but are not unimportant for the discussion on climate effects of changing trace gas concentrations as discussed in a later section. 0 3 emits only in a small fraction of the whole spectrum between 9 and 10 Iilll and around 14 Iilll. Its concentration varies byabout ± 30% over the globe.

Hane! [87J has made spectral measurements of the longwave radiation emerging from the top of the atmosphere from the satellite NIMBUS-5. These spectra provide an excellent survey on the interrelation between the emission of the earth's surface and the atmospheric gases (Fig. 15). The emission of the atmosphere measured from the ground in different locations in the same spectral interval is shown in Fig. 16. Beyond 15 Iilll the spectrum is entire!y determined by the rotational band of water vapor (compare Fig. 7). There it exists nearly no transparency between ground and space, except under extremely dry conditions or at elevated sites.

At LTE (local thermodynamic equilibrium) each volume element radiates isotropically into all directions. Therefore the radiation emitted by the system does not have the strong azimuth dependence as observed especially at large solar zenith angles for scattered solar radiation. Variations in azimuth are only observed if the water vapor and aerosol distribution is not horizontally homogenous [84]. The observed radiance depends therefore primarelyon the zenith angle of observation according to the radiating air mass [84].

The basic approach es to compute the distribution of far infrared radiances and long wave radiation fluxes are discussed in a later sec ti on.

Radiation Properties of Clouds

Cloud droplets are large Mie scattering particles. The interaction of the clouds with the radiation field can therefore be treated with the same theoretical methods as aerosols. Because of the large optical thickness of clouds in comparison to a hazy atmosphere it is absolutely mandatory to include multiple scattering in the computations. The large cloud droplets are especially efficient scattering centers for the near infrared. Under blue sky conditions the ratio of the radiances for wavelength of 2 Iilll and 0.5 Iilll is in the order of 1: 100 while under altocumulus this ratio can be reduced to 1: 6 [82]. The much smaller dependence on wave!engths of the scattering coefficient gives the cloud the white appearence as compared with the blue sky.

The strong foreward scattering characteristic of a single cloud droplet is smoothed by multiple scattering within the ensemble of cloud droplets. In the near infra red the clouds absorb radiant energy in the absorption bands of water at 1.5 and 1.9 Iilll, which are partly overlapping with the atmospheric water vapor bands at 1.38 and 1.87 Iilll. For an estimate of the absorption in water clouds it is therefore important to know how much the water vapor on top of the cloud already absorbs of the radiation which would otherwise be absorbed in the cloud.

F or the absorption of radiation in the cloud its total liquid water mass turned out to be a more critical parameter than the cloud drop size distribution [88-91].

184

TE ::l.-

-; ~

'" 7E u

~ CI> u c: c

'" c a:

H.-I. Bolle

:~ ~~' J o~ .~.IM&<>/~_V

7 8 9 10 ~-'il 12 13 1~ IS

10 H2 ·C·K

8 Z'N,·IC

6

~

2

o "00 " .. 7 8

2

12

S. Agoto. 24.8.1961. Nr. 4097

0558_0658 CET (0551i.Q656 TLT)

ZoO·

'00

13 TL IS

BEER SHEVA. 155.1963. NR.7038

12 38 _1401 GMT (14S8 -16 2I TLTl

Z = 0°

O~7 ------~8------'9C-----~I~O-------I T------~12~----C13~-----I-L-------TS

Wavelen gth (pm)

Fig. 16. The atmospheric zenith emission spectrum between 7 and 15 Ilm as measured in three locations. (After Bolle [84]). Top: Jungfraujoch, Switzerland, 3,570 m altitude, 16.2.1961. Middle: S. Agata sue due golfi, Italy, 24.8.1961, 05:56--06:56TLT. Bottom: Beer Sheva, Israel, 15.5.1963, 14:58-16:21 TLT

This mass is primarily concentrated in the larger cloud droplets (10-20 J.Ull). Also the material of the condensation nuclei is important for the absorption. For soot nuclei Korb [91J found a strong increase of the absorption over that of pure water clouds.

For the albedo of the cloud the drop size distribution is equally important since the smaller particles which are dominating the number density are essential for the interaction with short wavelengths. It could be shown [89J however, that liquid water content and mean cloud droplet radius are satisfactorely related to each other so that the total liquid water content may be used as single modelling parameter. The characteristics of clouds as measured over the territory of the USSR are compiled in Table 9.

Tab

le 9

. Clo

ud c

hara

cter

isti

cs m

easu

red

over

the

ter

rito

ry o

f the

USS

R. (

Aft

er D

ubro

vina

[92

])

A.

Mea

n li

quid

wat

er c

onte

nt, g

m -

3

Tem

p. r

ange

°C

Num

ber

of m

easu

rem

ents

M

ean

liqu

id w

ater

con

tent

S

tand

ard

devi

atio

n P

aram

eter

izat

ion

Wo

cons

tant

s e

B.

Pha

se c

ompo

siti

on

Tem

p. r

ange

°C

15-1

0

527 0.

25

0.22

0.

197

0.05

Per

cent

age

of p

hase

s L

iq u

id

Mix

ed

Cry

stal

s

C.

Hei

ght

dist

ribu

tion

o -2

83.6

14

.1

2.3

10-5

1,49

4 0.28

0.

23

0.21

7 0.

06

-4

-6

69.0

25

.9

5.1

5-D

3,68

3 0.26

0.

21

0.20

7 0.

055

-8

-10

53.7

35

.4

10.9

0--

5

7,16

4 0.21

0.

19

0.18

8 0.

033

Clo

ud t

ype

Low

(St

, Sc

) L

ow (

St,

Sc, N

s)

Med

ium

(A

s, A

c)

Lev

elof

loca

tion

km

0.

3-D

.6

1-1.

5 3-

4 m

b

1,00

0-80

0 80

0-55

0 T

hick

ness

km

0.

4 1

0.5

D.

Par

amet

eriz

atio

n fo

rmul

a

Per

cent

age

of c

ases

wit

h m

ean

liqu

id w

ater

con

tent

<W

: N

(W) =

10

0(1

-e(W

-"I/

W )

.

Mea

n li

quid

wat

er c

onte

nt:

W =

W oe

<w 0 ~ W

o +

e W

ater

con

tent

of S

t an

d S

c.:

WSt

,sc =

c h

exp

[P(T

-2

73

)]jT

g m

-3

c=0

.7 -1

.0·

1O-3

g K

m-4

!! = v

erti

cal e

xten

sion

of t

he c

loud

T

= M

ean

tem

pera

ture

in K

P

=O

.l K

-1

-12

-1

4

36.7

41

.8

21.5

-5

--1

0

5,31

0 0.17

0.

15

0.13

2 0.

033

Hig

h (C

i)

7-8

550-

300

2

-16

-1

8

22.7

39

.9

37.4

-10

--1

5

2,73

4 0.15

0.

13

0.12

3 0.

037

-20

-2

2

16.8

35

.2

48.0

-15

--2

0

806 0.

12

0.07

0.

099

0.03

3

-24

-2

6

10.3

30

.9

58.8

-20

--2

5

234 0.

09

0.04

0.

069

0.03

9 -28

-3

0

5.9

50.0

44

.1

i. ::::-.

§ 8- i ~ 1 S' g- t!

l ~ i I f -00 '"

186

'" u c: ;:

E '" c:

" .=

'" u c: c

E w

'0 t O ~or

10-1

10-2

10-1

20 Wove leng t h {fl }

10

Cloud th lckness

A 2m B IOm C 50 m

A

B

C

o D

Wave numbe r {cm-I }

H.-J. Bolle

5

Fig. 17. Albedo, transmittance and absorptance ofwater clouds in the far infrared. (After Yamamoto, et al. [100])


In addition the shape of the cloud is critical for its reflectivity. In earlier studies exclusively horizontally infinite clouds have been investigated. McKee and Cox [93] compared the scattering ofvisible radiation for horizontally infinite and cubic clouds, which is still a rough approximation for areal cloud. In the horizontally infinite case the shortwave directional reflectance approaches unity for a vertically thick cloud. It is about 85% for an optical depth of 50 which corresponds to 1 km vertical extension, ifthe liquid water density is 0.2 g m -3 and drops to 20-40%, depending on the observation angle, for an optical depth of ab out 4. For cubic clouds the reflectance is reduced by about 30% at the larger optical depths Wendling [94a] found similar results for striated clouds. Their albedo is always lower than that ofhorizontally infinite clouds. Harshvardhan et al. [94 b] computed the emissivity for cubic clouds and found that it is always less than that for a plane parallel cloud of the same optical depth.

In the near infrared the reflectance of the cloud is reduced by the absorption in the water bands. In a theoretical study Stephens [88] treated a number of cloud types and determined the albedo, absorption and transmission for the total shortwave range. For very thick clouds the asymptotic values of 80% albedo and 20% absorption are reached. The absorption, however, varies with the amount of water on top of the cloud and would be about 5% larger without any water vapor on top. For clouds in the planetary boundary layer the computed optical properties correspond reasonably weIl with the theory [89]. Observed albedos range between 0.49 and 0.75 (theory 0.52--0.68), transmissivities between 0.12 and 0.67 (theory 0.18--0.69), absorption 0.04--0.14 (theory 0.09--0.12). Equally good agreement between measurements and theory have been found by Feigelson et al. [95,96], see Table 10.

The highest shortwave albedo have cirrus layers on top of cumulonimbus clouds (86%). The theoretical treatment of these ice clouds is difficult because the ice crystals are not spherical [97-99].

Also in the far infra red a number of studies have been performed on water clouds ofwhich we quote Yamanoto et al. [100] results in Fig. 17. The reflectance of clouds is very small in the wavelength region 8-50 1Jlll. Even for an infinitely thick cloud it is only 5% at 8 lJlll and 3% for 5-50 1Jlll. Thick clouds are therefore often treated as black emitters which is not completely valid but the uncertainty in the cloud top temperature determination also enters the computation in actual cases. Thin clouds, in the order of 10 m thickness, are, however, already highly transparent in the far infrared.

Ice clouds which emit already strongly in the far infrared can sometimes hardly be detected in the visible. In satellite pictures it is often found that cirrus shows up in 10 lJlll images but can not be discriminated from surface features at short wavelengths.

For the solar global radiation under a partly cloudy or overcast sky different authors have derived approximative formulas. Schulze [78] investigated the montly sums of the relative radiant exposure H tel = H; (measured)j H ~, where H 00 = JSoCOS (0(t)dt represents the radiant exposure at the top of the atmosphere, and derived the following relation with respect to total cloudiness N (in tenths):

(124a)

188 H.-J. Bolle

Table 10. Comparison between eomputed and measured radiation quantities of clouds. (After Feigelson [95, 96])

A. Albedo and absorption in the water vapour bands

Wavelength 0.712--0.765 0.81--0.865 ranges, nm Bands a 0.8

{! CI: {! CI:

Ca1culated 0.53 0.07 0.58 0.05 Experimental 0.60 0.08 0.60 0.07

April 10, 1971 St

Ca1culated 0.82 0.10 0.83 0.08 Experimental 0.65 0.26 0.70 0.23

Sept. 24, 1972 As, Fr, Ns, St. fog

B. Total albedo and absorptanee

Date April 6, Oet.l, 1971 1972

Cloud type St Sc

Cloud top height in km 0.3 1.05 Cloud base height in km 0.20 0.75 Optical depth 5 15 Water vapor density in gjm3 10 5 Solar Zenith angle 39 74 Measured albedo 0.27 0.65 Ca1culated albedo 0.30 0.66 Measured absorptanee 0.01 0.11 Ca1culated absorptanee 0.07 0.07

and for daily sums:

0.89-1.0

{!U 'I:

{! CI:

0.47 0.13 0.56 0.14

0.71 0.23 0.65 0.30

April 10, 1971

St

0.85 0.40

20 10 26 0.59 0.56 0.05 0.12

1.085-1.21 1.28-1.585

tP

{!

0.40 0.46

Oet. 5, 1972

Sc

0.85 0.30

25 10 50 0.66 0.69 0.15 0.13

1p

CI: (!

0.33 0.29 0.42 0.45

Sept. 24, 1972

CI:

0.51 0.46

As, Fe, Fog Ns, St

3.90 2.0

90 5

54 0.71 0.77 0.20 0.15

H;el(N)=(1-0.5 N-0.3 N 10) H;el(N=O). (124b)

No distinction was made between cloud types. The numerical values for the ratio Hrel(N)/ H rel (N = 0) are compiled in Table 11.

Thompson [101] used a parabolic fit for a similar date set and found

H(N) I-B H(O) = 1+ (I_N)o.61 '

where B is an empirical constant. Davies et al. [102] considered clouds in different layers i of the atmosphere

and computed the downcoming irradiance by a summation over the different layers which contain c1ouds:

I

H·=H~ IIrN.i(I-QcNQJ (126) i= 1


Table 11. Values for H'd(N)jH,,}(O). (After Schulze [78])

N 0.0 0.2 0.4 0.6 0.8 1.0

~,:~~L{Eq. (9.50a) 1.0 0.89 0.77 0.64 0.46 0.20 H,,}(N-O) Eq.(9.50b) 1.0 0.90 0.80 0.70 0.57 0.20

Table 12. Constants of the empirical formula of Davies et al. [99] to compute H in Jm - 2

Cloud Fog Ns St Sc As Ac Cs Ci Cu Cb type

100a 64.48 46.87 99.61 145.3 163.3 219.8 364.6 344.0 b 0.028 - 0.167 0.159 0.104 0.063 0.112 0.148 0.079 12, 0.66 0.6 0.5 0.2 0.51 0.51

H 00 is the irradiance at the top of the atmosphere, Qc the reflectance at the cloud base Qs the surface albedo (0.13). The transmission functions are computed by means of the empirical formula

(127) where

(128)

is the air mass according to Kasten [71]. The constants a and bare given in Table 12.

Kasten and Czeplak [103] have recently published a study on the dependence of the radiation fluxes at the ground on cloudiness and solar elevation. They arrived at the conclusion that independently of the season the total short wave hourly mean irradiances under a cloudy sky may in the average be represented by

E(N) = {1-0.75 (~r.4} E(O), (129)

where N is the fraction of the cloud covered sky in octals, and E(O) the irradiance for the cloudless sky. The diffuse component Ed increases with N by

Ed(N) = {0.3+0.7 (~r} E(N). (130)

For cirrus clouds the ratio E(8)j E(O) for the completely overcast sky (N = 8) versus cIear sky is, however, 2.5-3 times as large as for low and medium clouds, for which it is in between 0.2-0.3. For nimbostratus this ratio is up to 30% smaller than for low and medium clouds.

Twomey [104] and Grassl [105] have investigated the problem that clouds may change their optical proper ti es due to the addition of air pollutants. A small effect to albedo increase can be expected if aerosols are imbedded in thin cIouds while for thick clouds the absorption due to the aerosols get more important and the albedo will be decreased.

190 H.-J. Bolle

Radiative Properties of Earth Surfaces

The radiative properties of earth surfaces exibit high variability, especially in the shortwave region of the spectrum. Here the reflection properties are important for the determination of the albedo. With only few exceptions their emission properties in the far infrared are much more uniform.

Water Surfaces

The most simple surface to deal with appears to be the water surface. But this is only true with respect to still clean and deep waters. Here it is possible to apply Fresnels reflection formula and to compute the reflection from the refractive index of water. The result for e.g. 812 nm wavelength is a specular reflection of about 2% for vertical incidence which increases to 3% at 50°, 10% at 66°, and 38% at 80° [106]. For an oceanic water surface the specular reflection is broadened into a sunglint pattern due to the surface waves, and is enhanced due to foam. Part of the light penetrates into the ocean and is scattered by small air bubles and by suspended particulate matter. Therefore a scattered sub-surface component is added to the upwelling radiance. The reflection function of ocean water depends on the wind speed by wh ich the water is stressed and which genera te waves, foam and air bubIes. The properties of the ocean surface are therefore also variable in space and time. The description of theses properties goes back to the pioneering work ofCox and Munk [107] and has been improved by other authors [108-111]. For illustration a selection ofvalues for the reflection function, which was defined in Table 13, of the ocean surface is presented in Fig. 18. Köpke [112] constructed a sea surface model with the sub-surface radiance data according to Table 13 and an isotropie foam albedo of 45% [113], whieh covers between 0% (up to 4 m S-1

wind speed) to 1 % of the surfaee area. At 7 m s - 1 wind speed the foam covers 0.45% of the surfaee [114].

The shortwave albedo of a sea surface as determined by different authors is given in Table 14.

Table 13. Values of the isotropie reflection function for the sub-surface contribution of the radiation reflected at water surfaces in percent of the incident radiation. (After Köpke [112])

90 Model with v = 1.4 m/s Model with v = 7 m/s A=0.55 I1m 0.7 !1ffi 0.55 11m 0.7 !1ffi

82S 0.28 0.077 0.38 0.107 77S 0.31 0.085 0.39 0.108 72S 0.34 0.094 0.41 0.110

67S 0.40 0.107 0.42 0.112 57S 0.44 0.113 0.45 0.117 47S 0.46 0.118 0.46 0.118

32S 0.45 0.108 0.46 0.114 12S 0.44 0.105 0.45 0.112 2.50 0.44 0.104 0.45 0.112


For the heating of the ocean the penetration of shortwave radiation is of importance. Smith and Baker [122] have recently reviewed available data of the attenuation coefficient in clear ocean waters from which the Fig. 19 was derived.

The emissivity of water at longer wavelengths is the parameter which determines together with the surface temperature the emission from the surface. For quiet water it can again be obtained by application of the Fresnel formulas

c 0

v c

" u..

c 0

u cu

Q; a:::

~ 0.1

Angle 01 rellection (degreel

Fig. 18. Reflection function of the ocean surface. (After Köpke [112])

192 H.-I. Bolle

Table 14. Sea surface albedo

Clear sky Overcast

( 0 =80° ( 0=20°

Payne [115] 0.28 0.035 0.061 Anderson [116] 0.21 0.04-D.05 0.04-D.05 Hollman [117] 0.03 0.05 Grischenko [118] 0.33-D.47 0.05 Kondratyev [119] 0.27-D.32 0.05

Computed

Burt [120] 0.24 0.05 0.06 Ter-Markariantz [121] 0.26 0.04 0.07

as difference between incident and reflected radiation flux. The reflectance spectrum of water in the far infrared and its variation with incidence angle is shown in Fig. 20. Between 4 and 13 j.Ull the emissivity of still waters is according to this Figure for large incidence angles 2± 1 %, but a destinct wavelength dependence is apparent. For ocean surfaces Köpke [125J uses a value of 0.979 devived from more recent data of the refractive index of water [126J for a zenith angle of 50° wh ich is in good agreement with the values presented in Fig. 20.

0.1

E

~ Range from var iou5 sources c: 2 1.0 0 :> c: ~ ~

..c ä. Q)

."

c: 10 -0

~ Q)

c: Q)

a..

100 200 800

Wavelength (n m]

Fig. 19. Penetration depth of shortwave radiation for clearest ocean waters. (After Smith and Baker [I 22 bJ)


Q) u c: ~ u Q)

't a:

Q) u c: '" Ü Q)

't a:

0.075--.------------------------~

0.050

0.025

2 3 4 5 6 7 8 9 10 11 12 13

Wavelength in)-lm

0.5 --.---------------------------,

0.4

0.3

0.2

0.1

2 3 4 5 6 7 8 9 10 11 12 13

Wavelength in}-lm

193

Fig. 20. Reflectance spectrum of water in the far infrared and its variation with angle of incidence, after [123 a, b, 129] .

194 H.-J. Bolle

Solid Earth and Vegetation Surfaces

The albedo of the solid earth is determined by the mineral composition of the soil and by its vegetation cover.

The following major reflection features can be distinguished: 1. The angular distribution of the reflected radiation depends on the structure of the surface. There exist surfaces which reflect the radiation nearly isotropically into all directions. Others have a specular component, and a third category exibit astronger backward scattering component. Coulson [127, 128] has investigated the reflection of desert sand, in comparison to dark clay (see also [86]). It is found that the desert sand scatters more radiation in the backward direction, that is in the direction of the illuminating source (the sun), than in the foreward or specular direction. Dark clay on the contrary has its reflection maximum in the direction where the specular reflection peak can be expected. In the sand the radiation can penetrate into the uppermost layers due to the grannular structure. Here it is scattered according to the multiple Mie scattering law but the primarely strong foreward component is dispersed. The clay, because of its very fine dustlike particles, behaves more like a mirror, especially for longer wavelengths. Also asphalt as used for street pavements exhibit this strong foreward peak.

Plants, beeause oftheir irregular shapes, tend to scatter the light more isotropic. Here the shadowing effect becomes, however, also important: vertically protuding vegetation may have stronger backward reflection from its illuminated leafs than specular reflection because an ob server from this direction looks into the shadowed areas [131, 132].

The angular distribution of the radiation scattered at the surface is. very complicated and has therefore to be studied for each type of surface separately.

2. The spectral distribution of the reflected shortwave radiation depends on the chemical composition of the surface and its humidity. Soils reflect generally less at shorter wavelength than in the near infrared, which is partly due to the grain size effect. Shorter wavelength can easier be absorbed in cavities between grains than longer wavelengths. The spectral reflection of soil materials is generally a smooth curve. Moisture reduces the reflection due to absorption.

The spectral reflection properties of plants are domina ted by the chlorophyll absorption at 0.4 and 0.6 /lIll. Therefore in the visible part of the spectrum the reflection is generally low, depending on the kindand color of the leaves, and around 0.7 /lIll there is a sud den steep increase to reflectances higher than 50%.

In most of the reflection spectra the near infra red absorption bands of liquid water show up, and sometimes also the water vapor bands are not completely removed.

3. With increasing wavelengths the structure of the reflection spectra in the longwave infrared get more complex because here absorption bands of the individual minerals become essential [130].

The fundamental vibrations of the mineral lattices in the far infrared act as resonators for the incident radiation which is therefore strongly reflected if no other absorbing mechanisms (like surface cavities) are present. Such reflection bands are called "Reststrahlen" or residual bands (such material were and are used for spectral filtering by multiple reflection in the far infrared). One of the


'~ Ul

06

0.5

0.4

0.2

0.1

-- Pasture land ......... Bog

--- Savannah

1

1 I I I I 1

OSun

Backward scatteri ng

I .•..........•. 1 ,-,

1 " \ I "

0.87pm

0.87fm 1 , ,

1 /' \,/ I I .... 1

1

' f',.... #

-'- Coniterous torest

Os2pm

... 1./ -7"" ...... I / ./ /" . I I ,... // ",., ...... ,.............. /-, i;' ... f:;...::::.~;,·/

/' 'r' ... ········· / '. ...,,/' 1./ /'

~ '._- ... ,-_ -""r .0°1.

'''~--- \.......................... I/I ."-., I " /1

0 ......... _.-"-"" I . . I '" .... \ 1..-/

'<~-~·····;:~2·~·~·~~;,~·:~:::::·:~·_·_// Zenith angle ot observation

Azimuth = 00 Azimuth = 1800

195

Fig. 21. Bi-directional reflectance of pasture land, coniferous forest, savannah and bog. (After Kriebel [133])

196 H.-J. Bolle

most important of these features is the Si02 band at 9 J..Ull [106] which is also apparent in the spectrum taken over the Sahara in Fig. 15.

Humidity may decrease the reflectance by a factor of 3 near 2 J..Ull. For four surfaces, pasture land, bog, savannah and coniferous forest, Kriebel

[133] has determined the bidirectional reflectance at wavelengths for different solar aspectand observation angles. Ofhis data two sets are selected to demonstrate the main features of the four surfaces (Fig.21). These are data for 40° incident angle and 0° respectively 1800 reflectance azimuth with respect to the direction of incidence. The reflectance of the vegetated surfaces is less than 0.03 at 0.52 J..Ull for most of the observation angles but rises at large zenith angles. Savannah, due to its sparse vegetation, has a slightly higher reflectance. In the backward scattering direction a slight peak is observable. At 0.87 J..Ull the reflectance of pa sture land is strongly enhanced due to the high near infrared reflectivity of fresh gras. Coniferous forests are much darker because they act as light-traps. Savannah and bog lie inbetween. Savannah exhibits a strong backward reflection peak. Because of its isolated trees and bushes the reflection at specular angles is suppressed. The same is true but due to another reason for coniferous forest, which acts as a light trap because of its pyramidal tree structure. The bog behaves more like pasture land but is much darker.

Of interest in this respect is a thorough investigation carried out by Romanova [134] on the reflectance of sands on the territory of the USSR. She found that the reflectance of sand barchans and quartz sand increases between 500 and 600 nm, that the reflectance of sand barchans does not depend on the direction of the sand ripples with respect to the solar aspect angle and that polymietic sands (consisting of different sediment components) and graywacke (strong feldspar components) have smaller reflectance; graywacke has also a smoother wavelength dependence.

Rott [135] has determined the reflectance of an urban area from aircraft and Landsat images (Table 15) which shows a rather uniform urban reflectance throughout the visible part of the spectrum and the chlorophyll features in the rural area (grass land).

The use of satellite data to determine albedo of extended areas gains in importance. However, careful corrections for the atmospheric attenuation and scattered solar radiation have to be applied. From the measured directional

Table 15. Directional reflectance of urban and rural area determined from Landsat and Aircraft images. (After Rott [135])

Aircraft (solar zenith angle 35°)

Wavelength in ~ 465 515 560 600 640 680 720 815 Rural reflectance in % 5.6 6.4 7.2 6.3 5.7 5.4 14.9 28.5 Urban reflectance in % 7.3 7.2 7.2 7.5 7.6 7.3 8.5 11.4

Landsat (solar zenith angle 39°)

Wavelength in ~m 500-{)00 600-700 700-800 Rural reflectance in % 6.7 5.7 21.7 Urban reflectance in % 7.2 7.5 10.0


reflectance albedo values can only be derived, if the reflection function, the angular distribution of the reflected radiation is known [136].

A thorough discussion of spectral albedo was recently published by Kondratyev [139]. A compilation given by Kondratyev earlier [119J is reproduced in Tables 16 and 17. More data can be found in [68, 137].

The emission at wavelengths longer then 2.5 Jlill depends as weIl on the surface structure and its grain sizes as on the refractive indices of its materials, and the soil humidity. Where the surface reflects radiation its emittance will be less than one, reflectance and emittance add to unity. Most surfaces and especially those which are wet can approximately be treated as black bodies. The exception is quartz sand which has its reflection band directly in the atmospheric window where it is most effective as can be seen in Fig. 15. The emissivity of different natural surfaces has been measured by Buettner and Kern [138J, see Table 18.

Snow and I ce Surfaces

The directional reflectance and albedo of snow varies largely depending on the freshness of the snow deposition, its liquid water content and structure. Its spectral variability throughout the visible part of the spectrum is small, but throughout the near infrared there is a decrease until after a steep step near 2.7 Jlill the radiation is nearly completely absorbed. In the reflection spectrum ice absorption bands can easily be detected. They are present at 1.5 11m and 2.0 11m and are thus slightly shifted to longer wavelengths with respect to the water vapor and liquid water bands. For ice surfaces the optical effects depend on the ice structure. In the visible ice is dark if it has no cracks, but enclosed air bubbles and structural inhomogenities act as scattering centers as weIl.

Wiscombe and Warren [140J developed a detailed model for snow albedo. They treated the snow particles as Mie scattering centers and applied radiative transfer theory (Delta-Eddington approximation, compare following section on computation methods). The dependence ofthe albedo on grain size and imbedded impurities was investigated in detail. At wavelength A< 1 Jlill the theory results in a steady decrease of the albedo of pure (deep) snow from 0.98 at 0.4 Jlill to about 0.70 at 1 Jlill with still 0.95 at 0.7 Jlill. In the infrared the albedo is modulated by the ice absorption bands. At 1.5 Jlill wavelength the albedo drops below 0.05 with revivals for new snow at 1.85 Jlill and 2.3 Jlill (albedo 0.15). In nature even at remote stations like the South Pole [141J the maximum of the observed spectral albedo for aged snow occurs around 0.9 Jlill. The depression at shorter wavelengths is interpreted as aerosol pollution. Actually measured spectral albedo values can be modelled by imbedding desert dust or soot particles into the snow model.

Wagner [142J measured the development of the albedo during aperiod when freshly fallen snow was transformed by a melting period at the equilibrium li ne of a glacier. Directly after a snowfall of more than 5 cm precipitation in July the albedo of the fresh snow was 95% (diurnal mean value). Over the wet glacial ice finally only 15-20% albedo were measured. Other authors [143, 144J found albedo values down to 6% over ice. According to Wagner the albedo measured during the ablation period had the record which is summarized in Table 19.

Tab

le 1

6. S

pect

ral

alb

edo

of

nat

ura

l su

rfac

es u

nd

er a

cle

ar s

ky.

(Aft

er K

ondr

atye

v [1

19])

Sur

face

h

o,

Wav

elen

gth,

nm

'" 00

deg

400

450

500

550

600

650

700

750

800

850

900

950

1,00

0

Soi

ls a

nd

ro

ad c

over

s

Wh

ite

san

d (

rive

r)

47

34.0

37

.0

40.0

44

.0

46.0

48

.0

49.0

50

.0

54.0

A

sph

alt

road

48

13

.2

16.5

18

.2

19.2

19

.8

20.4

20

.9

22.0

22

.0

24.0

C

on

cret

e ro

ad

47

12.4

13

.6

15.3

15

.1

16.6

20

.4

22.3

22

.8

22.9

23

.2

25.0

C

on

cret

e in

the

fo

rm

50

13.0

22

.0

29.0

34

.0

38.0

39

.0

40.0

40

.0

40.0

o

f li

ght

slab

s D

irt

road

(dr

y)

50

8.0

10.0

12

.0

13.0

15

.0

20.0

22

.0

24.0

26

.0

26.0

D

irt

road

(w

et)

50

7.0

8.8

9.0

9.6

12.0

16

.8

19.4

20

.8

21.6

21

.6

Veg

etat

ive

cove

rs

Su

dan

gra

ss

52

2.0

2.8

2.8

5.0

3.8

3.0

6.0

43.0

50

.0

51.0

51

.0

51.0

A

lfal

fa (

J une

) 56

1.

8 2.

0 3.

5 6.

5 5.

2 4.

8 11

.7

35.0

40

.0

42.0

42

.0

42.0

41

.0

Cab

bag

e 50

5.

0 5.

8 6.

8 8.

0 9.

4 7.

6 13

.0

37.0

42

.0

44.0

44

.0

41.0

G

reen

thi

ck g

rass

56

2.

0 2.

8 3.

6 5.

2 6.

2 5.

8 5.

6 23

.0

39.6

40

.0

39.2

38

.5

37.7

A

lfal

fa (

J ul

y)

53

1.8

2.0

2.8

4.2

3.5

6.0

9.2

24.0

28

.0

30.0

30

.0

30.0

30

.0

Clo

ver

55

1.

8 2.

0 2.

4 4.

2 4.

4 3.

8 8.

0 23

.0

30.0

32

.0

33.0

34

.0

Th

in g

rass

50

3.

0 5.

6 7.

0 5.

0 4.

5 8.

4 19

.0

27.0

28

.0

29.0

29

.0

30.0

B

ienn

ial

vine

yard

50

2.

3 2.

9 3.

3 5.

0 5.

2 5.

0 14

.0

22.0

25

.0

27.0

28

.0

29.0

29

.0

Sil

age

corn

54

2.

0 2.

3 3.

0 5.

8 5.

4 4.

2 8.

0 28

.0

32.0

33

.0

34.0

35

.0

35.0

T

all

gree

n m

aize

56

3.

8 5.

0 7.

6 7.

6 7.

3 10

.8

24.0

29

.0

28.0

29

.0

30.0

30

.0

Yel

low

co

rn

46

5.0

7.0

8.0

10.6

11

.8

13.0

16

.4

22.4

24

.0

25.0

27

.0

28.0

29

.0

Sun

flow

er

52

1.5

2.0

2.5

7.9

7.1

6.5

10.0

21

.8

25.0

27

.0

28.0

29

.0

29.0

S

tubb

le o

f ce

real

s (t

hick

) 35

6.

0 7.

6 8.

2 10

.2

12.9

15

.5

18.6

20

.3

28.0

30

.5

32.5

33

.2

33.9

S

traw

56

6.

0 8.

0 11

.0

14.4

20

.6

26.0

30

.0

36.0

38

.0

40.0

43

.0

43.0

M

ow

ed s

ilag

e co

rn

52

4.2

5.6

7.0

8.0

7.6

7.4

9.0

14.0

15

.2

15.8

19

.9

20.4

21

.1

Dry

gra

ss

28

3.5

4.5

5.6

7.0

8.7

11.0

13

.6

16.5

19

.3

21.0

22

.5

23.5

24

.0

See

dlin

g o

f w

inte

r w

hea

t 48

4.

0 4.

3 5.

0 6.

6 7.

2 7.

2 10

.4

14.0

17

.2

19.4

21

.0

21.8

22

.6

Soi

l af

ter

thaw

ing

of

sno

w

24

2.0

2.5

4.0

5.8

7.0

6.8

11.0

15

.5

16.7

17

.8

20.1

20

.8

20.8

~

Sn

ow

cov

er

, :-<

Dry

fre

shly

fal

len

sno

w

30

85

87

90

90

90

90

90

90

89

87

86

81

78

I:l:l

0

Wet

fre

shly

fal

len

sno

w

30

83

85

87

87

87

86

86

86

86

84

82

76

70

~

Wet

sn

ow

38

62

66

67

67

69

68

67

65

62

58

51

46

42


Table 17. Dependence of albedo from solar zenith angle for some of the surfaces listed in Table 16. (After Kondratyev [119])

Surface Solar zenith angle in degree

40 45 50 55 60 65 70 72

Chernozem 1.00 1.00 1.02 1.04 1.06 1.08 1.10 1.13 Sparse grass 1.00 1.00 1.03 1.08 1.16 1.22 1.24 1.27 Green thick grass 1.00 1.03 1.07 1.14 1.19 1.25 1.29 1.32 Silage corn 1.00 1.04 1.07 1.12 1.21 1.32 1.42 1.45 Yellow corn 1.00 1.05 1.09 1.20 1.37 1.54 1.69 1.76 Mowed silage corn 1.00 1.00 1.07 1.15 1.33 1.59 1.89 2.00

Table 18. Vertical emissivity of various materials measured by Buettner and Kern [138] and other sources

Material EIT Material

Quartz (agate) 0.712 Human skin Granite 0.815 Sand, quartz large grain Feldspar 0.870 Sand, quartz large grain wet with Obsidian 0.862 water (nearly saturated) Basalt 0.904 Sand, Monterey, quartz small grain Dunite 0.856 Concrete walkway, dry Granite, rough side 0.898 Asphalt paving Obsidian, rough side, 0.837 Water, pure

broken glass appearance Water, plus thin film of Basalt, rough side, shiny 0.934 petroleum oil Dunite, rough side 0.892 Water, covered by a thin sheet of Silicon sandstone, polished side 0.909 polyethylene Silicon sands tone, rough side 0.935 Melting snow Dolomite, polished side 0.929 !ce Dolomite, rough side 0.958 Vegetation Pla te silicon glass 0.865 Sahara

Table 19. Albedo at the equilibrium line of a Glacier (Hintereisferner, 2960 m, Austria, 1971). (After Wagner [142])

Surface

Fresh now >5 cm, Ta;,< 273 K ::::5 cm, Ta;,~273 K

Aged snow, pure Aged snow, polluted Neve, pure Neve, polluted lee mixed with water, poilu ted

Albedo in %

90-95 68-90 62-68 46-59 (50-65) 25-30 15-20

EIT

0.980 0.914 0.936

0.928 0.966 0.956 0.993 0.972

0.961

0.995 0.98

>0.97 0.8-0.9

200

1000

500

300 200

100

'I 50 '-1Il

, 30 E 20 :::L

N

'E 0 3

5 (lJ u

3 c 0

2 -0 0

0:: 0

J 5

0 3

02

o I

- ."- ' - ' - ' -.

05 06

......

0.7 0.8 0 .9 1 .0

= } . . . . }

- ' - '

1.2

H.-J. Bolle

snow coverM

Iree 01 SnOw

1.4 1.6 1.8 2.0 2 .2 2.1.

Wavelength (p m)

Fig.22. Radiance from snow covered and snow free areas observed from Jungfraujoch. (After Bolle [145]). ----11.30 TLT, -16.00 TLT, ····14.00 TLT, -. -·16.00 TLT.

The effect of a snow covered glacier surface in comparison with a snow free area in spring can clearly be seen in Fig. 22 where also the absorption bands near 1.5 and 2.0!illl and the chlorophyll absorption edge at 0.7!illl stand out clearly. The radiance ratio snow covered to snow free is about 7 in the visible.

Basis for the Theoretical Treatment of Radiative Transfer

General Remarks

One of the ultimate goals of the theoretical analysis is an improved quantitative understanding of the interaction between the radiation field, aerosols and clouds in order to be able to compute with high accuracy the energy fluxes in the earthatmosphere system, and to determine the distribution of absorbed solar energy between the atmosphere, the land surface, and the ocean. To meet the requirements


for an accurate treatment of the shortwave and longwave radiation field in the atmosphere it is necessary to apply rigorous radiation transfer theory. For extensive application e. g. in numerical climate models the computation procedures as described below are often much too elaborate and computer-time consuming. Therefore more simple but nevertheless accurate "parameterized" computation procedures need to be developed as weIl. Such computations need accurate input data on atmospheric and surface parameters. For simplified approaches the empirical evidence is an excellent guide, and comprehensive da ta sets are needed to extract the statistical properties ofthe system. A large amount of such observational data has e.g. been compiled by Kondratyev [119].

In this section a short outline is given from what basic principles such computational procedures are derived, and a list of references is provided where the specific methods are discussed in detail.

Radiative transfer is more generally treated in a number of monographs of which those concerned with radiative transfer in scattering atmospheres were already quoted in an earlier section [11-16]. For longwave radiative transfer the starting point - the general radiative transfer equation - is the same as in the case of shortwave radiation but the numerical procedures are different. For shortwave radiation the modelling of the highly anisotropic scattering processes is the main problem. For the terrestrial emission the treatment of the various atmospheric spectral bands require special efforts in either computing the cumulative effect of all single lines or in deriving appropriate parameterisations, so-called band models. The latter will however, not be discussed in detail here. The reader is referred to the comprehensive monographs of Goody [39], Möller [146], Kondratyev [147], and Paltridge and Platt [148].

7he General Equation of Radiative Transfer

If a geometrical path along the direction s is considered at the height z in the atmosphere, the radiance can be denoted by L(z,' s) in a horizontally homogenous atmosphere. The direction s can be expressed by the zenith angle ( and the azimuth <po As vertical coordinate the optical depth of the atmosphere, i5= SO"edz, will be used. The increment of the geometrical path, d.), will be replaced

by ds = ~, where /l = cos (. The radiance L(i5; <p, /l) is attenuated due to extinction between sand 0" e/l

s + ds by the amount di5 {dL(i5; <p, /l)}ex, = - L(i5; <p, /l) --.

/l (131)

It will on the other hand be increased due to scatter from other directions s' into the direction s and/or by emission according to

• di5 { dL(b; <p, /l)} add = J(b; <p, /l; <p', /l') O"eJ1 (132)

where J (b; <p, /l; <p', /l')= 0"; 1 J (b; <P, /l; <p', /l') is the sourcefunction for this additional radiance. The total change in L is the sum of Eqs. (133) and (132) which can be written after division by db//l:

(133)

Equation (133) is the general radiative transfer equation (RTE), which applies to all radiation problems. The sign of the l.h.s. depends on the choice of the positive direction for s.

202 H.-J. Bolle

1he Source Function Jor a Scattering Atmosphere

The source function can be derived from an inspection of the involved process. For the moment only scattering and absorption in aerosols will be considered. In this case the incremental radiance emerging from a small volume in the direction <P, Jl results from scattering of radiation which traverses the volume in all other directions. Consider a direction <p', Jl' in which the radiance L(o; <p', Jl') proceeds within a solid angle dQ'. This radiance generates a scattered radiance dL(O; <P, Jl; <p', Jl') in the direction <p, Jl within the incremental solid angle dQ of the magnitude

(134)

Here pro; <P, Jl; <p', Jl') is the scattering phase function defined in Eq. (27) which prescribes the distribution ofthe scattered radiation uL(o; <P, Jl') to the different directions, dQ'/4n is the ratio of the incremental solid angle to the sphere from where the radiation comes and dQ is the small solid angle into which the radiation is scattered. The geometrical path ds in which the scattering occurs is again expressed by do/ueJ.I.. According to Eq. (5b)u/ue =w isthe single scattering albedo,and for the scattering phase function the normalization

...!...J (<- .' ')drl - -- ~<1 4 pu,<p,Jl,<P,Jl .&-w- = n U e

(135)

applies. In order to account for all directions from where radiation arrives in the scattering volume,

Eq. (134) has to be integrated over all solid angles dQ' which resuIts in

{ dL(o; <P, Jl)} W ( J L«· , ') (<- .' ') dQ') Jl do sc = 4n 4nQo v, <P , Jl . pu, <P, Jl, <P , Jl , (136)

By comparison with Eq. (133) one can see that the r.h.s. of equation (136) is the source function J for this process.

It can be seen that the solution of the radiative transfer equation is complicated by the scattering properties of aerosols.

Especially if large particles contribute to the scattered radiance the computation of the integral in (136) is difticult to perform because of the strong forward peak in the scattered radiance.

Methods to Compute Radiances and Fluxes in a Thrbid Atmosphere

In order to get the radiance at any point and in all directions, Eq. (133) has further to be integrated over the whole atmosphere and all higher orders of scattering have to be regarded, incIuding reflection at the ground.

There exist different methods to solve the RTE which have recently be reviewed by Lenoble et al. [149, 150], and which can only shortly be described and referenced here.

Monte Carlo Method [151-155]

In this method the path of one photon after another is computed assuming a random chain of scattering and/or absorption processes. After a large number (e.g. 106) of photons has been tracked the number of photons arriving at the detector are counted and set into relation to the initial number.

Approximation by Spherical harmonics [156-158]

The phase function is first separated in azimuth by expansion in Fourier series mostly by the use of Legrendre polynomials. Then the radiation transfer equation is expressed in Legendre functions and a System of first order linear differential equations results for the coefficients of the expansion which has to be solved.

Method of Discrete Ordinates [159-162]

The equation of radiative transfer is separated in azimuth and the integrals replaced by sums. A set offirst-order nonhomogeneous differential equation is generated which can be solved numerically.


Gauss-Seidel Iteration [163-166]

The atmosphere is divided into layers and the unit sphere into a set of solid angle inerements. Thus the optieal depth and he solid angle are diseretized and the radiative transfer equation is approximated by a set of linear algebraie equations.

M ethod of Successive Orders of Scattering [167-171]

The intensity is expressed as aseries developed for the number of seattering proeesses to be eonsidered. Then the equation of radiative transfer is solved after separation in azimuth.

Discrete Space 1heory (Matrix Operator Method) [172-176]

The laws for transmission, refleetion and absorption (interaetion prineiple) are formulated for diserete layers of the atmosphere. Refleetivity and transmissivity are represented by matrix operators. From the formulation for one layer the method is extended to two and more layers by applying matrix multiplieation formalism.

Doubling or Adding Method [177-184]

The refleetion and transmission of eaeh of two layers are eonsidered as known. Then the refleetion and transmission of the two layers together ean be obtained by eomputing the baek-and-forth refleetions between the two layers. If the initial thin layers are of equal depth the result for a thiek layer ean be build up by geometrieally adding two layers of equal depth.

Invariance Principle [184, 185]

This is another method where like in the ease of Monte Carlo eomputations no direet attempt is made to solve the radiation transfer equation. A number of logical relations are defined whieh re1ate upwards and downwards direeted irradianees by means of generally defined transmission and refleetion funetions to eaeh other. These are the invariant relations, whieh are differentiated for the optieal depth and the derivatives are eliminated by means of the radiative transfer equation whieh enters as a side relation. The generated system of equations is solved by expanding the seattering indieatrix whieh enteres with the radiation transfer equation by means of Legendre polynomials.

Invariant Imbedding Method [186-190]

The invariant imbedding method starts from a eritique of the solution of the RTE by means of expansions of the phase funetion and linearization of the equations. If the seattering is highly anisotropie - as in the ease of large aerosols - the linearized algebraie systems are ill-eonditioned and give instable or only slowly eonvergent solutions. In the invariant imbedding method the integro-differential equations for the transmission and refleetion funetions are direedy treated. The integrals are approximated by sums via Gaussian quadrature. The resulting ordinary differential eq uations are then integrated numerieally.

Dodecaton Approach to Radiative Transfer (DART) [191]

In this method spaee is subdivided into a eertain number (whieh ean be arbitrarily large) of radiation streams arranged on a regular dodeeahedron. The minimum number of diserete streams whieh has to be eonsidered for apre-set aeeuraey ean be determined from Bayes' Rule of probability theory.

Approximative Similarity Solution [192]

The radiation transfer in a anisotropie seattering medium is treated as an isotropie seattering problem in whieh the single seattering albedo and the optieal depth are adjusted by means of an asymmetry faetor.

Eddington Approximation [193, 194]

In the Eddington approximation the field of seattered radiation is approximated by an isotropie distribution of the radianee whieh depends on the optieal depth of the atmosphere and the solar zenith angle only. The phase funetion is approximated by p(9) = 1 + ßI eos9, where ßI is eonstant.

Schuster-Schwarzschild Approximation [195-199]

In this approximation the ineident flux is represented by a parallel beam whieh is broken up due to seattering in one eomponent parallel with the ineident direetion and one in the opposite direetion. The phase funetion is thus represented by a Delta-funetion.

90°

204 H.-J. Bolle

60o-r---

300

'" Öl c 0° 0

.c

c '" N

~ 30°

60°

90°-1----

180° 180°

Fig. 23. Distribution of downwelling (Jeft) and upwelling seattered solar radianee in a turbid atmosphere eomputed by Bakan and Quenzel [204]

Exponential Kernel Approximation [200,201]

The phase funetion is approximated by a Delta-funetion and two Legendre terms. The exponential integrals appearing in thc formal solution of the equation of radiative transfer are approximated by exponential funetions in which a diffusity faetor of 3/2 appears in the exponent.

Pertubation Method [202-204]

In this method only variations of the single scattering albedo with optieal depth are eonsidered. On a referenee homogeneous atmosphere of eonstant single albedo a variable perturbation is superimposed and the deviation from the referenee ease is eomputed.

These more elaborate eomputations of the radianee field in scattering atmospheres result in distribution for the c1oud-free atmosphere Iike the example reprodueed in Fig. 23. There are hardly any direet eomparisons between eomputed and measured radianees because the simultanous determination of all necessary atmospheric parameters and of the speetral radianees is diffieult to aeeomplish. As far as eomparisons could be made the eomputed radianees compare favourably with the features measured from the ground.

90°


Top of Q t m 0 5 P her e ~r-r-----------------------------------------,O

N

~ z L tJ)

<lJ

I

Transmission layer for dL

Geometrical path s

Zenith angle

d u

Do

O~ ~ Earth surface

Fig. 24. Derivation of the equation of radiative transfer for an emitting atmosphere

Radiative Transfer Equation for the Terrestrial Emission

205

A computation of the radiance distribution has to start from Eq. (132) which defines the source function. If an inclined path is taken in a plane parallel atmosphere (Fig. 24) the contribution of an atmospheric layer of the geometrical thickness dz in the direction ( is

dL=amQ(z) B[T(z)] dz /l

(137)

where am is the mass absorption coefficient and Q(z) the density of the emitting gas. Only the fraction

{ ' dz} dz (dL),~o=(dL),r(/l; z, O)=am(z)Q(z)B[T(z)] exp - SamQ(z)- -o /l /l

(138)

arrives at the surface z=O because of absorption between z=O and z. With the optical depth between z and z = 0, ,

6(z, 0) = S amQ(z)dz,

the transmittance is r(O, z) = exp { - 6(z, 0)/ /l}. Equation (138) can then be transformed to

o

_ (dL)FO= B[T(z)] dr(/l~zz, 0)

(139)

(140)

206

and the emission of the wh oie atmosphere becomes

00 d,(W z 0) L(Il, z=O)= - fB[T(z)J d ' dz.

o z

A more general RTE can be derived from Eq. (133) in which the source function replaced by B [T(z)J defined in Eq. (88):

dL'(O,Il) ") (")J 11 do L,{u - B,[T u .

H.-J. Bolle

(141)

J has to be

(142)

Here the spectral index 11 has been added. To compute the irradiance, Eq. ([42) has to be solved for L,(o, 11) and to be integrated over all directions 11. A formal solution [see Eq. (39)] is

E,= f L"dQ=nfB,[T(O,,)J d{2Ei3 (0,,)}, 2n:Qo

where Ei3 is the third exponential integral 00

Ei 3 (u)= f e- ux x- 3 dx 1

(143)

which can be approximated by a simple exponential function (ie-'U) with a r-times elongated optical depth u. The irradiance can therefore be approximated by

E, ~n f B,[T(o,)J d,,(r· 0).

r is called diffusity factor and has a mean value of 1.66.

Parameterization of the Atmospheric Band Structures

(144)

To determine the total irradiance Eq. (144) has to be integrated over all wavelengths. The difficulty to solve the RTE (144) lies in the determination of the transmission functions. Several gases with highly variable absorption coefficients and partly overlapping bands contribute to the radiation flux. Different approach es are in use to deal with this problem.

a) ComputationJrom Basic Line Parameters [84,205, 206J

The atmosphere is subdivided in a number of levels with constant temperature and density of the absorbing molecular species (H2 0, CO2 , 0 3 , CH4 , N2 0). For each wavenumber the transmission from each level to the observer (e.g. ground or space) is computed by setting a equal to the sum of all individual line contributions [Eq. (77)]. In the 8-13 J.Ull window region also the water vapor continuum and the aerosol contribution has to be added as described earlier. The radiance is then computed by replacing the integral by a sum over the atmospheric layers:

N

L,= I Bb ;[1] Ll i " (145) where

i= 1

Lli'='(Zi' 0)-,(Zi-1, 0).

In order to reproduce a spectrum with high accuracy the computations have to be done in steps of 10- 3 - 10 - 4 cm - 1 in the vicinity of line centers. The atmosphere has normally to be subdivided into 30--100 layers. One example of a computation is given in Fig. 7 [84].

b) ComputationJrom Averaged Empirical Band Transmissivities

The spectrum is subdivided according to the contributions of the major molecular bands and the slope of the Planck function. For each spectral band empirical transmission functions or averaged transmission functions computed from basic line parameters and averaged Planck functions are used. Transmission functions have been measured for specific temperatures and a range of press ures [45-49]. In order to account for the temperature and pressure dependence of the transmission often standard transmission functions are used but the masses are scaled by (p/Po) (Ta/T)n [207]. The computations then follow the procedure of equation (145) with all quantities averaged over the spectral


intervals. Finally the partial speetrally averaged radianees are summed up to eompose the total radianee.

It is also possible to average first the transmittanee speetrally for all absorbers and to solve Eq. (145) only onee. This ean then be done graphieally by means ofradiation diagrams [208-210].

e) Emissivity Method [211, 212]

The RTE ean be re-formulated by introducing the emissivity instead of the transmission. The emissivity 8 at pressure p and temperature Tfor the mass m for the whole speetrum is defined by

1 <X>

8(m, p, T)= -T4 f a,,(m, p, T) B,,(T) dv. (J 0

With 1 -, = 8 the irradianee ean thus be written

E= JB,,(z') (d8,,(Z, z')) dz' z dz

, z· <X> nB,.(z") da" _ " 8,,(Z, z)= f J uT4(z") dz" dvdz .

(146)

(47)

The integrals over the wavenumber ean be computed in advance so that for the direet application only the integration over the atmosphere remains to be solved.

General Energy Budget Equations for an Earth-Atmosphere System

The global energetics deve10p from an interplay of the processes which occur on regional scales in areas of varying boundary conditions. It seems appropriate therefore to start the discussion with a description of the energy processes in a cell or a "cage" which extends from the ocean or solid earth to the top of the atmosphere. The cage will extend into the ocean or solid earth as far as it is affected by energy transports. For the oceans this depth is difficult to assess since climate variations may affect the deep ocean but only after a long time period. The exchange between the deep ocean and its upper layers is in the order of several hundred years. If we restrict ourse1ves to shorter time periods the deep ocean will be decoupled from the system and we need only to be concerned about the upper layers of the ocean, the "warm water sphere" or the top layer of the ocean with a few hundred meters thickness (compare Fig. 1).

If the flux divergencies in such a cage or box can be determined from its physical properties and chemical composition then it willlater on be possible to connect all these boxes in order to construct a global system [213,214].

To establish the energy budget of the volume all energy fluxes into and out of such an e1ementary box of the earth-atmosphere system have to be considered (Fig. 25). The externat input is the solar energy flux which may be partly compensated or in other locations overcompensated by the infrared emission to space. It is possible that at the surface some of the solar energy is stored for some time, generally half a year, if the system is in a steady state. This implies that there is no accumulation or loss of energy over periods of one or more years.

There are other energy transfers through the vertical walls of the volume, transports which are accomplished by the general circulation of the atmosphere and the oceans. These fluxes are called advective fluxes. It is assumed that horizontal radiative transports are neglegible and that only heat as well as potential and

208

<P~w,~ =

RpSoG (Ce)

r---------------~ I I I LiQA= i I * I I <PR(z) agas,ae Li!: I I I I I I L ______________ ....J

r----I

t <PLW,~

Cloud

<P;(z) ac Li! = LiQc

H.-J. Bolle

I===::t> <PA,adv,o

<Po, adv, i ====::::t>F=========t>l: F===========t=====t> <Po, adv, ° I I I I L ____________________ ~

Fig. 25. Energy fluxes within and across the boundaries of an ocean-atmosphere volume

kinetic energies are transported this way. These transports compensate for any gain or loss of radiant energy within the box over longer time scales. Excess energy of one volume is horizontally transported to areas with energy deficits. The mechanism set into action to maintain these transports is the general circulation of the atmosphere and the oceans.

Within an elementary volume energy conversions as discussed in the introduction take place. At the surface energy arriving as solar radiation is transformed to conductive heat fluxes into the surface and the atmosphere. A substantial part of the energy is used to evaporate water which is transported as latent energy into the atmosphere. The heat flow into the atmosphere at the ground generates buyancy which carries heat upward by turbulent processes. Latent heat can be released if condensation occurs and clouds are building up. The heat released during condensation warms the atmosphere further and potential energy builds up by the rising air masses, generating pressure gradients which eventually require an outflow of these elevated masses across the walls of the volume, compensated by an inflow at lower levels. Thus potential energy disappears and kinetic energy is genera ted, which finally is destroyed by friction at the surface and inbetween the air molecules.


In order to quantify the relations between the different fluxes of energy in the earth-atmosphere system fundamental mathematical relations are sought now. The imaginated volume extends a few hundred meters into the oceans or a few meters into the soi!, as far as the annual winter-to-summer temperature variations reach. In Fig. 25 only an ocean is considered for the lower boundary of the atmosphere. In the case of solid earth there would be no horizontal fluxes through this part of the volume but the picture would essentially remain the same.

At the top ofthe atmosphere the solar irradiance So is incident under a zenith angle (0' It generates a flux <P§w,00=SOAG((0) into the volume. G((0) is a geometryfactorwhich would beidentical to cos (0 for a plane parallel atmosphere and deviates from this function slowly if (0) 75°. A is the horizontal area of the atmosphere-surface system. Part of the shortwave radiation is reflected and scattered back into space. This fraction is (lp <P§w, 00 where

(148)

is the planetary albedo. The atmosphere-earth-surface system will also emit infrared radiation,

<pLw, 00, to space. The radiation budget at the top of the atmosphere is therefore

<Pl oo = <P§w, 00 (1- (l) - <PLw, 00' (149)

The budget is positive if energy is provided to the system. The star is used for the difference of the fluxes with opposite direction through a surface and is referred to as net flux.

At the surface a fraction of <P§w,oo arrives partly as direct [<P§w, 00 r(O, 00)J

and partly as scattered or diffuse solar radiation <P~,o. The sum of these two components is called global radiation <P~,o:

<P~,o=SoAG((dr(O, 00)+<P~,0. (150)

The radiation budget at the surface consists of four components, the downwelling shortwave and longwave fluxes, the reflected shortwave flux and the flux of the emitted longwave radiation:

<P~, 0 = <P~, 0(1- (}o) + <pLw, 0 - <PLw, o· (151)

(lo is the surface albedo. From the surface the fluxes of sensible and latent heat are directed into the

atmosphere, and another heat flux into the ground, the ocean or the soi!. In these media the energy may be stored or transported away. The energy budget of the ground volume is therefore

* * _ iJQGR _ 1 iJTGR <PR,o+LI<Po,adv-~H,o-<PLH,O-<PPH- ~- (}c ---at. (152)

LI<PO,adv=<PO,adv,j-<PO,adv,O> influx minus outflux through ocean,and <PPH is the flux which is used for photosynthesis and which is therefore stored in biological mass (less than 1 % of <P§w,o). That part of the net flux which is stored in the ground, <PbR, 0' raises the temperature of the surface layer of the ground and thus affects the emitted longwave radiation which will be raised according to the

210 H.-J. Bolle

surface temperature until the opposite process starts. The heating respectively cooling which occurs is expressed by the r.h.s. ofEq. (152).

The budget ofthe atmosphere is in analogy to Eq. (152):

,M m* m* Am _ 8Q A _ 1 8 TA 'l'R oo-'l'R o+'l'H+LJ'l' A adv- -8-- --8

" 't {!Cp t (153)

Here the latent and the sensible heat fluxes are combined to cI>H and the star is used in order to indicate that there are also reversed fluxes due to dew and precipitation.

Combination of Eqs. (152) and (153) results in

~; = cI>~, 00 + L1 cI> adv - cI>PH, (154)

where 8Q/8t is the rate of storage in the whole system. If cI>PH is neglected in Eq. (154) and if the fluxes are averaged over a year or

a number ofyears one can assurne that 8Q/8t=0 (if there are no climate changes), and therefore

<P~, 00 = L1 cI> adv' (155)

Averaged over aperiod of one year (or several years) the radiation budget at the top ofthe atrnosphere equals the divergence of energy advected into the volurne by both, the oceans and the atrnosphere.

If Eq. (155) is furthermore zonally averaged it is clear that there can be no flux across the poles and Eq. (155) can be rewritten in the following way. The earth is subdivided into N latidudinal zones i, defined by corresponding geographicallatitudes cP i'

[L1cI>advJ;= [<Padv(CP)] - [<Padv(CPi - dJ. (156)

Applied to Eq. (155) it results

N N

L [<P~, 001= L [L1 cI>adv] i = <Padv(CPN)- <Padv(CPo)' i= 1 i= 1

(157)

If the summation starts at one pole (CPo=900) where cI>adv(CPo)=O, then the sum of N zonal mean net radiation fluxes at the top of the atmosphere represents the advective flux across the latitude CPN'

The zonally averaged advective fluxes in the earth-atmosphere system can therefore be determined by careful measurements of the zonal radiation budgets from space. This quantity is a boundary condition for the processes in the system but does not yet provide much insight into the processes within the system itself, especially because it is not yet possible to discriminate between oceanic and atmospheric transports and to define those transports which finally determine the environmental or internal system conditions.

From satellite observations and computations using conventional data [215-219] it can be concluded that the total transport is in the order of the numbers given in Table 20.

In order to gain more insight into the functioning of the system the next step is to inspect more closely wh at happens at the borders of the box volume and

I"

10

Tab

le 2

0. Z

onal

ly a

vera

ged

radi

atio

n bu

dget

par

amet

ers

and

mer

idio

nal f

luxe

s. (

Aft

er E

llis

and

Van

der

Haa

r [2

38]

but

bala

nced

with

res

pect

to

tota

l ene

rgy

budg

et)

~

~.

0

Lat

itud

e zo

ne

Are

a 4

nR

2 A

rea

in

Len

gth

of

Ann

ual

mea

n B

alan

ced

zona

l H

oriz

onta

l fl

ux

Hor

izon

tal

flux

:;

10

in d

egre

e in

%

1012

m2

(10°

)-1

zona

l ci

rcle

so

lar

irra

dian

ce

net

ener

gy i

nput

ac

ross

lat

itud

e pe

r un

it l

atit

ude

:;

p..

in 1

06m

in

W m

-2

in P

J (1

0°) -

I s -

I in

PW

in

GW

m-

1 m

:;

'" .... 0<>

'<

80-9

0 N

0.

760

3.88

17

3.9

-0.4

05

1 ...,

6.94

8 +

0.40

5 +

0.05

8 ....

70-8

0 2.

255

11.5

0 18

5.5

-1.0

87

~ '"

13.6

88

+ 1

.492

+

0.10

9 '0

60

-70

3.68

5 18

.79

213.

5 -1

.37

3

0

20.0

12

+2.

865

+0.

143

::l

."

5'

50-6

0 5.

000

25.4

9 26

0.5

-1.2

14

....

25.7

27

+4.

079

+0.

159

'" ~ :;.

40

-50

6.16

0 31

.41

307.

0 -0

.68

6

~ '"

30.6

58

+4.

765

+0.

155

m

30-4

0 7.

140

36.4

1 34

7.7

-0.0

08

0

10

34.6

61

+4.

773

+0.

138

Z

.... 20

-30

7.90

0 40

.28

380.

1 +

0.69

6 :;.

37

.610

+

4.07

7 +

0.10

8 g

10-2

0 8.

420

42.9

3 40

2.4

+ 1

.914

39

.415

+

2.16

3 +

0.05

5 0

0-10

8.

680

44.2

6 41

4.1

+2.

566

'" '0

40.0

24

-0.4

03

-0

.01

0

po

0-10

S

8.68

0 44

.26

414.

1 +

2.44

2 '" ....

39.4

15

-2.8

45

-0

.07

2

'" CIl 10

-20

8.42

0 42

.93

402.

4 +

1.7

08

'<

37.6

10

-4.5

53

-0

.12

1

~ 20

-30

7.90

0 40

.28

380.

1 +

0.84

9 S

34.6

61

-5.4

02

-0

.15

6

."

30-4

0 7.

140

36.4

1 34

7.7

-0.0

16

.... '"

30.6

58

-5.3

86

-0

.17

6

~

40-5

0 6.

160

31.4

1 30

7.0

-0.8

87

..<

: 25

.727

-4

.49

9

-0.1

75

'5

50

-60

5.00

0 25

.49

260.

5 -1

.48

7

0

20.0

12

-3.0

12

-0

.15

1

CIl

60-7

0 3.

685

18.7

9 21

3.5

-1.6

26

j 13

.688

-1

.38

6

-0.1

01

70

-80

2.25

5 11

.50

185.

5 -1

.04

1

6.94

8 -0

.34

5

-0.0

50

80

-90

0.76

0 3.

88

173.

9 -0

.34

5

Tot

al

100.

0 0.

5099

. 10

15 m

2 0.

0

~

.....

212 H.-J. Bolle

what transformations occur within the volume. We start with this discussion at the top of the atmosphere.

Energy Fluxes at the Top of the Atmosphere

Solar Irradiance

The upper boundary of the atmosphere is distinguished by the fact, that energy transfer occurs only by radiation, of which the solar radiation is the single incident component. Its normal incidence flux area-density, the solar constant So, has been determined from high altitude observatories by extrapolation of spectral measurements to zero air mass [61, 62J, as weIl as from aircraft [63, 220J, Balloons [65, 221-224J, rockets [225J and more recently also from satellites [226-229]. The conclusions drawn from two different subsets of the data presented in Fig. 26 differ considerably. Vonder Haar et al. [230J and Fröhlich [231J arrive at an unweighted mean of 1,377±20Wm-z with a standard deviation of 8Wm- z. Brusa and Fröhlich [232aJ and Crommelynck [232bJ give a value of 1,367±5Wm-z. The higher values are not in accordance with other radiation budget studies [229J, where a value of 1,368 W m - Z would be preferable. A conservative figure for the solar constant would therefore presently be

So=I,370±8Wm- z . (158)

The solar constant was believed to be about 9 W m - Z lower than this value more then a decade aga and rose to more than 1,370 W m - Z in the early 1970 ies, but these discrepancies can be explained by the improving but still not perfect measuring techniques. There is therefore no indication of a real trend in this fundamental quantity [232c]. Only Kondratyev and Nikolsky [222J reported unconfirmed changes of So in the order of 2.5% over aperiod of five years. Strong

Solar Solar Solar Solar Minimum Maximum Minimum Maximum

1410 3 5 9

j 7

':'E 1390 - j 1 ,/,1 A 13

3

f''' Nimbus-6< Nlmbus-7

A 14 ~18 ~ f /17 ~ 1370-

" J Rocket __ •

t2 c 1976 15 Rocket t t 0 1978 u 16 19 20 21 ~ 2 4 6 MarinerVI,Vll ~ 1350f-lf)

1330 1964 1 65 I 66 1 67 1 68 1 69 1 70 1 71 1 72 1 73 1 74 1 75 1 76 1 77 1 78 1 79 I 80

Ye a r

Fig. 26. Experimental detenninations of the solar constant during the last decennia. (After Fröhlich [231], Vonder Haar et al. [230], and Brusa and Fröhlich [232])


variations which increase in amplitude with decreasing wavelengths below 200 nm are observed in the uv part of the solar spectrum [233]. Brusa et al. has recently also reported [232a] short term fluctuation with periods between 5 and 20 min of the solar flux. The amplitude of these short-term fluctuations are in the order of 0.1%0, again growing with decreasing wavelength. Willson et al. [232d] detected variations of the total solar irradiance from the Solar Maximum Mission spacecraft around the mean value of 1,368.31 W m- 2 in the order of ±0.05%. Such fluctuations occur in the frequency range below 0.15 day-l. Two larger decreases in irradiance of up to 0.15%, each lasting for about one week, have also been observed. These two decreases are apparently correlated with the development of large sunspot groups.

A new attempt to measure the solar constant will be made from future space missions. By means of the carefully designed instruments which are nowadays available, it should be possible, to measure the long term variability of the solar constant to 0.1 % or 1.4 W m -2 precision, and to correlate measured variations with the solar rotation or solar activity. The absolute mean value stated in Eq. (158) may in the future still be object to corrections by a few W m -2.

The knowledge of the solar constant is of fundamental importance for the determination of the equilibrium state of the earth. Variations of the mean solar input will immediately induce climate variations.

The geographical distribution of the solar irradiance at the top of the atmosphere is determined by astronomical parameters like the earth orbit around the sun and the inclination of the earth rotation axis with respect to the plane of the earth orbit, the ecliptic.

The earth orbit is an ellipse with a ratio minor axis to major axis of 0.967. The mean distance between sun and earth (=semimajor axis of orbit) is called 1 Astronomical Unit (A.u.) for which the internationally accepted value is 149.5.106 km. The earth presently is at its mean distance from the sun on April 3 and October 5. The minimum distance (147.1 .106 km), the perihelion, is passed on January 5 and the maximum distance (152.1 . 106 km) the aphelion, on July 5.

The rotation axis of the earth is tilted against the orbital plane by 23°27'08':2 at the present time. Thus the sun is in the equatorial plane of the earth only twice a year, on March 21 and September 23, at equinox. The north pole is tilted against the sun on December 21 and the south pole on June 21. These positions are called solstices. The angle between the equatorial plane and the sun is the declination <5 0 which accordingly varies during the year between about + 23~5 (lune 21) and -23~5 (December 21).

The irradiance at the position of the earth depends on the distance between sun and earth. The geographical distribution of the irradiance on top of the atmosphere is a function of the solar zenith angle (0 which can be expressed by the geographical latitude qJ, the declination <5 0 , and the time between local noon and the ho ur of observation. Since one rotation of the earth around its axis takes 24 h for 360°, the angle of rotation corresponding to one ho ur is 15°. Therefore

h=(15· Lltot, (159)

where h is the angle of rotation corresponding to the difference Ll to between the true local time (TL T) t and true local no on in hours:

Llto = t-12.00 (true local time). (160)

Ll to is negative in the morning, positive in the afternoon. The True Local Time (TL T) is related to the Local Mean Time (LMT) by the equation oftime, Ll't, which is normally presented in tabular form (72):

Ll't=TLT-LMT. (161)

214 H.-J. Bolle

This difference arises because the rotation of the earth is not constant during the year. The variation has the magnitude of -14.33 min (12. Feb.) to + 16.4 min (3. Nov.) with two secondary peaks ( + 3.73 min on 13. May and - 6.42 min on 27. July). Clocks are measuring Zone Mean Times (ZMT) which generally change by one hour every 15° longitude starting at Greenwich (0°). This has to be taken into account if one transfers ZMT to TLT:

t(TLT)=(A-ARl·';' + t(ZMT) + L'l't , (162) where AR = ZMT reference longitude A = longitude of observation site in degree (positive to east). From spherical geometry the following equation for the solar zenith angle can be derived if the reference plane is tangential to the earth surface:

(163)

The irradiance at the top of the atmosphere is

_ S=So(~r COS(0. (164)

So is the irradiance at d= 1 A.u., and d the actual distance between earth and sun. The integration of Eq. (164) over the day gives the radiant exposure measured in Ws m - 2 or

Jm- 2 per24h:

( (5) 8.64.104 (d)2 [h· ." " . h] (165) Hcp,0=--n--Sod smcpsmu0+cosCPcosu0sm.

8.64 . 104 are the seconds per day and h is the length of the half day, e.g. from sunrise to no on in radian and is defined by Eq. (163) with( 0=0:

cosh = - tg cp tgi5 0 .

For locations and times where the sun does not set it is h = n and

H(cp, i5d=8.64· 104 so(~r sincp sini50 ·

For i5 0 = +23.5" this happens to be for

1 tgcp= tg23.5 =2.3, or cp=66~5N.

(166)

(167)

As an example the equator at equinox (i5 0 =0, cp=O, d/d~l) receive 37.6· MJ m- 2 per day. At solstice the (north) summer pole (i5 0 = 23~5, cp= 90°, d/d=0.967) receives by the factor (d/d)2 h sin 23.5 =0.9672 . n . sin 23.5 = 1.17 more energy than the equator at equinox; and the equator (cp = 0) receives at solstice 0.9672 cos 23.5=0.857 times the energy incident at the equator at the time of the equinox.

PIanetary Albedo

A fraction of the solar radiation is scattered or reflected back to space. The ratio of the outgoing shortwave flux density M to the incoming solar irradiance S is the instanteaneous planetary albedo Qp of the earth at the geographical coordinates cp and A:

(168)

where S((0) is given by Eq. (164). The planetary albedo in a certain area is a function of the cloudiness and the

albedo ofthe clouds, the turbidity ofthe atmosphere, the scattering characteristics of the aerosol, and of the surface albedo. It is quite difficult to compute the


planetary albedo since multiple scattering processes in the atmosphere-surface system have to be taken into account. If one disregards this complexity, one may approximately write

Qe, h, Qe, m' Qe, J

Qs

Qa

N=Nh+Nm+NJ

r(O, 00)

are the fractions of coverage for high, medium and low cloud tops in the measured area the albedos of the high, medium and low clouds respectively surface albedo the albedo of the cloudless atmosphere due to moleeule and aerosol scattering total cloudiness transmission of the cloudfree atmosphere from the ground to space

In Eq. (169) the first three terms on the r.h.s. include the albedo of the atmospheric gases on top of the clouds.

The albedo produced by the scattering atmosphere, Qa, depends on the concentration of aerosols in the atmosphere and its scattering characteristics. The molecular component of the atmosphere pro duces an average albedo of approximately 8% depending on the mean solar zenith angle (Table 21) which has amaximum value of39S at equinox for the equator. Va lues ofthe planetary albedo are compiled in Table 20 after computations made by Braslau and Dave [234, 235]. If the effect of molecules, aerosol, water vapor, ozone, carbon dioxide and oxygen is considered, the average backscatter of the cloudless atmosphere is about 10%, a value which is also observed over the oceans outside the solar reflex [236].

The albedo of clouds depends on the type of cloud, water cloud or ice cloud, and the thickness or liquid water respectively ice content of the cloud. There have been several attempts to determine the average planetary albedo of a cloudy sky and its geographical distribution either by computations based upon cloud observations, local surface albedo measurements and aerosol estimates, or by direct measurements made from satellites. In Fig. 27 some of the more recent results from satellite measurements are reproduced.

The mean value of the planetary albedo which is presently the most probable one [229, 238] is

Qp=0.305. (170)

The range of earlier determination was between 0.35 [218] and 0.29 [215].

Terrestrial Longwave Radiation

Infrared radiation is emitted by the earth surface, cloud tops and atmospheric gases like water vapor, carbon dioxide, ozone and, to a much lesser extent, CH4

and nitrogen oxides. The total emission from a black surface into the hemisphere is according to the law of Stefan and Boltzmann given by

(171)

Tab

le 2

1. P

lane

tary

alb

edo

for

a m

id-l

atit

ude

sum

mer

atm

osph

ere

afte

r B

rasl

au a

nd

Dav

e [2

34,

235]

Sol

ar z

enit

h an

gle

Alb

edo

of u

nder

lyin

g su

rfac

e

Ray

leig

h sc

atte

ring

R

ayle

igh

scat

teri

ng a

nd

gas

abs

orpt

ion

Ray

leig

h sc

atte

ring

, gas

abs

orpt

ion

and

a)

av

erag

e ha

ze L

[26

] ae

roso

l sc

atte

ring

b)

ave

rage

aer

osol

sca

tter

ing

and

abso

rpti

on

Ray

leig

h sc

atte

ring

, gas

abs

orpt

ion

and

a)

heav

y ha

ze L

[26

] ae

roso

l sc

atte

ring

(4 t

imes

inc

reas

ed p

arti

cle

conc

entr

atio

ns i

n th

e st

rato

sphe

re a

nd

in t

he b

ound

ary

laye

r)

b) h

eavy

aer

osol

sca

tter

ing

and

abs

orpt

ion

ALBE

DO

ANNU

AL

ZONA

L M

EAN

S 9

00N

o

0°

0.0

0.1

0.05

0.

14

0.04

0.

12

0.05

0.

12

0.05

0.07

0.

14

0.06

• E

MIT

TE

D

MIN

. E

) E

MIT

TE

D

MA

X .

• A

LBE

DO

M

IN.

o A

LBE

DO

M

AX

. I!;

, M

EA

N

0.3

0.27

0.

26

0.26

0.27

90

.51

W·

....

0.1

0.

2 0

.3

04

0

.5

0.6

0

.7

Fig.

27.

Ann

ual a

vera

ge a

nd v

aria

bili

ty o

fzon

ally

ave

rage

d pl

anet

ary

albe

do d

istr

ibut

ion

with

lat

itud

e.

(Aft

er C

ampb

ell

and

Von

der

Haa

r [2

37])

30°

0.0

0.06

0.

05

0.06

0.

06

0.09

0.08

60°

0.1

0.3

0.0

0.1

0.3

0.14

0.

33

0.09

0.

18

0.35

0.

12

0.26

0.

07

0.14

0.

28

0.13

0.

27

0.10

0.

16

0.29

0.

09

0.15

0.

29

0.16

0.

22

0.28

0.14

80°

0.0

0.1

0.17

0.

25

0.15

0.

21

0.22

0.

27

0.21

0.34

0.

38

0.31

0.3

0.41

0.

32

0.37

0.45

IV

0\ :x::

., ~

o:J o ~


where (J= 5.6696.10- 8 W m- 2 K- 4

is the Stefan-Boltzmann constant, and T the temperature in degree Kelvin. The relation (171) results from an integration of Eq. (86) over all wavelengths.

The maximum ofthe emission for terrestrial temperatures per unit wavelength interval is around 10 J.Ul1 (for unit wavenumber interval around 600 cm -1 or 17 J.Ul1 compare Figs. 15 and 16) following from Planck's law, Eqs. (86), (87), which is graphically presented in Fig. 8. In the region between 8 and 13 J.Ul1 wavelength the atmosphere is rather transparent so that a direct exchange between surface and space can take place, except for the region of the ozone band (9-10 J.Ul1). The water vapor continuum as weIl as aerosol absorption do, as mentioned earlier, weaken this direct radiant energy transfer. The emission of a specific area of the earth can be computed from the contributions of the different constituents of the emitting system. If again the N j , i=l, m, or h, denote the fraction of low, medium and high cloud coverage as seen from the top of the atmosphere, and T</>(Z, 00) being the spectrally averaged transmittance for the flux between the altitude z and the top of the atmosphere, then the flux on top of area A will be with N = LN;:

j

3

c[>Lw=csMB(,T.) (l-N)T<1>(O, 00)+ L cjN jM B(TJT<1>(Zj, 00) j; 1

(172)

Cs and Cj are the emissivities of the surface and of the cloud tops respectively. The first r.h.s. term is the emission of the surface to space in cloud-free areas, the second term the contribution of three cloud layers, the third and forth terms the contribution of the atmosphere on top of the cloud free surface and on top of the three c10ud layers respectively.

Maximum temperatures occur at the surface in deserts or semi-arid areas. The highest temperatures are about 330 K. The lowest temperature ever recorded at the surface is 184.9 K (Antarctica, 78° S, 3,420 m altitude) and minimum temperatures occur in the tropical tropopause level with 180 K (compare Fig. 1). Temperature in the earth atmosphere system vary therefore by about 255±75 K or ±30%. Now the infrared emission varies with T 4 and if we take the extremes the maximum variability of local emissions can in fact be 1: 10. However, since the coldest temperatures are in the tropopause in lower latitudes where at the surface the warmest temperatures occur and since at the poles the cold surface temperatures are partly compensated by a relatively warmer atmosphere, the emission at the top of the atmosphere appear much more uniform. In fact the equator to pole gradient of the infrared exitance is only 2.2 in the annual mean for the southern hemisphere (compare Fig. 28).

A computation of this quantitiy needs considerable effort since the atmospheric temperature structure, the heights and emissivities of the clouds and the cloud distribution as seen from the top of the atmosphere have to be known. Satellites do in fact contribute essentially to determine some of these parameters

218

90 0 N

o

EMITTED FLUX ANNUAL ZONAL MEANS

• EMITTED MIN. EI EM ITTED MAX . • ALBEDO MIN. o ALBEDO MAX. 8. MEAN

H.-J. Bolle

900S~~~ 100 150 200

W/m 2 250 300

Fig. 28. Annual average and variability of zonally averaged longwave exitance of the earth-atmosphere system. (After Campbell and Vonder Haar [237])

[239, 240]. By means of special instrumentation and application of sophisticated evaluation procedures to operational meteorological satellite data the total outgoing longwave flux is also directly measured from satellites. One example of the results obtained so far is reproduced in Fig. 28.

Equilibrium Condition

For the limited time period of a few years it can be assumed that there is virtually no climate trend. Over periods of one or more full years the radiation budget at the top of the atmosphere must then be assumed to be balanced: there is no accumulation or loss of energy in the total system. Under these circumstances the absorbed solar energy must completely be re-emitted as infrared radiation. Let rE be the radius of the planet. nr~ is its cross section with respect to the parallely incident solar radiation, 4nr~ its surface, and Qp its mean planetary albedo. The absorbed solar energy is nr~So(1- Qp) and the emitted infrared radiation can be set equal to 4nr~CTT,,4 where T" is an equivalent emission temperature for the longwave radiation emerging from the different sources in the atmosphere: surface, clouds, aerosols and gases. The condition for radiative equilibrium is

For T" it follows therefore:

1:= e

(173)

4

(174)

With Qp=0.31 and So=1,370Wm- 2 the equivalent temperature T" becomes 254 K. For a standard mean atmosphere with a surface temperature of 288 K the temperature of 254 K occurs at about 5.2 km altitude. This level can therefore


be assumed as the average emission level, which results from the distribution of clouds and emitting gases in the atmosphere.

It can easily be seen what would happen, if either the solar constant or the albedo of the planet would change. Differentiation of Eq. (174) gives:

oTe loSo (175) Te 4 So

respectively oTe 1 oQ

(176) --~-

Te 41-Q

If the solar constant varies by 1 % then the equivalent temperature changes by 0.01 Te/4 or 0.64 K. In this estimate it is assumed that the albedo does not change during the transition which is not for sure. If on the other hand the solar constant remains unchanged but the planetary albedo changes by + 1 % absolute (which means e.g. from 31 % to 32%) then the equivalent temperature decreases by 0.92 K. In both cases the climate on earth will be distorted. To what degree the change in the equivalent temperature may be reflected in the surface temperatures depends on the structure of the atmosphere and its re-adjustment to the new conditions.

Equation (173) suggests that the total infra red flux area density

M I - 7"'4 - 1 S (1 ) - 1 E* LW, co -a 1 e - 4 0 -Qp - 4 SW,co (177)

is one forth of the net solar irradiance at the top of the atmosphere. The mean infrared flux density should therefore be 235.8 W m - 2 which has not yet been verified empirically to better than about 5%. The net radiation flux as measured from satellites is presented in Fig. 29 (units: Wm - 2).

Fig. 29. Annually averaged net radiation flux area density at the top of the atmosphere. (After Vonder Haar et al. [230))

220 H.-J. Bolle

Energy Fluxes at the Earth Surface

Radiation Budget

The radiation budget at the surface is determined by five components: the fluxes of the direct solar radiation, <p(], the diffuse shortwave radiation, <P~, and the infrared emission of the atmosphere, <Pt, are the downwelling components; the fluxes of the reflected shortwave radiation <P~ and the thermal emission of the surface <Pl,J are the upwelling components. The net radiation can be expressed by

(178)

where <Pt 0 = <PI. 0 - <Pl,t, and <P~ = <P (] + <P~ the "global radiation", (Jo is the surface albedo.

On a world wide basis so far only Budyko [241, 242] has determined the geographical distribution of the net radiation flux densities, and London and Sasamori [218] have determined the zonal mean net radiation flux densities at the surface. Perry and Walker [243] have given an excellent survey on the radiation budget over the oceans. The annual mean insolation at the earth surface varies between about 8 GJ m - 2 a - 1 in the tropics to a minimum of 2.5 GJ m - 2 a - 1 at high latitudes 8. The poles receive up to about 25 MJ m - 2 d -1 during summer, which is comparable to the daily radiant exposure in the tropics throughout the year. Due to the high albedo of the polar ice caps only a small fraction of this energy is, however, absorbed at the surface. Thus the annual mean absorbed shortwave radiation ranges between 6.5 GJ m - 2 a -1 in the tropics and only about 0.7 GJ m - 2 a - 1 at the poles [244, 245]. The infrared emission of the atmosphere depends on atmospheric temperatures and its water vapor concentration. It is maximum in the humid and hot zones of the tropics with 12 GJ m - 2 a - 1 and decreases to 4.8 GJ m - 2 a - 1 near the poles, which is still about 40% of the tropical value. The emission of the surface on the otherhanddecreasesby46%from 14.2GJ m- 2a -1 in thetropics to 6.6 GJ m- 2 a- 1

at the poles. The longwave net radiation is - 2 GJ m - 2 a - 1 in the tropics, -1.8 GJ m - 2 a - 1 around 65° S latitude, -1 GJ m - 2 a -1 over Antarctica [245], and - 0.85 GJ m - 2 a - 1 over central arctic [244].

In the southern hemisphere the zonally averaged annual mean radiation budget gets negative near 70° S. In the northern hemisphere it seems to remain positive

8 The absolute maximum values for the shortwave radiant exposures disregarding the effects of the atmosphere are those at the top of the atmosphere. The extraterrestrial radiant exposure at 1 A. U. in the tropics is for a Solar Constant of 1,370 W m -2 equal to 1,370 J m -2 per 1 s. For different time intervals the following numbers can serve as an orientation: 1 s at noon: 1.370 kJ m- 2 S-l

1 min at noon: 82.2 kJ m- 2 min- 1

1 hat noon (3,600 s): 4.93 MJ m -2 h- 1

1 day, computed after Eq. (165): 37.7 MJ m- 2 d- 1

1 month, sum over 30 days 1.13 GJ m - 2 month - 1

1 y, sum over 12 months 13.6 GJ m - 2 a - 1.

Thus the natural units to measure radiant exposures are per time interval :;::; 10 min :;::; 1 week :;::; 1 y unit kJm- 2 MJm- 2 GJm- 2


with0.47 GJ m- 2 a- 1 [218], thoughalso smaH negative values have been reported [245]. Thus, with the exception of Antartica where values down to - 0.3 GJ m - 2 a - 1 are measured [246], the zonal means of the annual radiation budget at the surface are everywhere positive with a maximum value of about 4.5 GJ m - 2 a - 1 in the tropics.

The energy absorbed by the surface has to be dispersed by transport processes which will be discussed in the following sections. Before entering this topic the variability of the radiation budget has to be discussed.

a) Seasonal Variations

At the equator the seasonal variations are small, the radiation budget remains close to 400 MJ m - 2 month - 1. The belts of maximum radiation balance (about 500-580 MJ m - 2 month - 1) shift from 20° S in Winter to 25° N in Summer. At mid latitudes (45°) the variations are between zero in winter to about 330 MJ m - 2

month - 1 in north summer over the continents and 420 MJ m - 2 month - 1 in south summer over the oceans. Most dramatic changes are found in Antartica at coastal stations where south summer radiation budgets can be as high as 500 MJ m - 2 month -1 which drop in winter to -100 MJ m - 2 month -1 (the corresponding numbers for inland stations are + 33 and - 50 MJ m - 2 month - 1

respectively). Variation of about 400 MJ m - 2 month - 1 between winter and summer occur

at the coast of southern California and in the north Carribbean, as weH as at the coasts of Chile, Argentinia, South Africa, south and east Australia and the Yellow Sea area.

b) Geographical Variations

Over the oceans, especially in the tropics, the radiation budget is always larger than over land because of the smaller albedo of the oceans and the high er mean temperatures of the land surfaces. The difference is on the average approximately 38% but vanishes at high latitudes. Maximum va lues of the radiation budget up to 21 MJ m - 2 d - 1 are found in summer in cloud free regions of the subtropical oceans (15-35°) such as the northern Arabian Sea and over subtropicallakes and reservoirs. Minimum values of - 130 MJ m - 2 month - 1 occur near the poles during winter. Relatively low values ( < 12 MJ m - 2 d - 1) can as weH occur at high latitudes (> 50°) with long snow and ice cover as in the warrnest regions of the earth where albedo approach es 40%, and where the emission is also high because of the high daytime surface temperatures.

c) Diurnal Variations

Under clear sky conditions the radiation budget takes its maximum value around local no on when the insolation is maximum (compare Fig.40). It drops to zero about one hour before sunset and remains negative during the night until one hour after sunrise. The magnitude of negative values depends on the humidity of the atmosphere, which determines the infrared emission of the sky. Typical clear sky nightime values range around - 100 to - 300 kJ m - 2 h - 1. The larger negative

222 H.-J. Bolle

values occur if a re1atively warm surface emits into a dry sky. These conditions are observed at highlands over snowcovered terrain where values down to -370kJm- 2 h- 1 have been observed [247, 142].

During daytime the budget is mainly determined by the insolation and surface albedo. Peak values of 2 MJ m - 2 h - 1 are typical for mid latitude summer days.

Clouds modulate strongly these during daytime otherwise nearly sinousodial variations. The negative values at night are nearly constant. Under den se clouds the radiation budget can drop to zero both at day and night.

Partitioning of Radiant Energy

If all the radiant energy which is absorbed at the earth surface would be used to heat the uppermost layer of the soi!, then the surface temperature would rise until there is equilibrium between the incoming solar flux and the outgoing infrared flux. For an albedo of 0.2 and a solar irradiance of 0.5So=680Wm- 2 the equilibrium would be reached for T s=313 K if the emissivity is 1.0. Such high temperatures are rarely observed in mid latitudes for which the assumed amount of irradiance is not unusual. In fact, the absorbed energy is immediately partitioned in sensible heat fluxes into the ground (ocean or soil) and into the atmosphere, a flux of latent energy into the atmosphere, and a small fraction which is used for photosynthesis:

(179)

In this equation the heat flux from the interior of the earth is al ready neglected (it would have to be subtracted from wGR,o) since it is very small compared with the other fluxes (approximately 0.06 W m - 2).

The eminently important question for the environmental c1imate is now, how large the fractions of these different fluxes are, and what factors are responsible for the magnitude of the individual fluxes.

The heat flux into the ground determines the heating of the soil and the oceans. The heat capacity of the ground material is responsible for the storage of this energy. The heat can be released again at times when the incident radiation flux is small thus reducing the rate at which the surface cools at night or during winter. The magnitude of latent flux determines the amount of water released into the atmosphere which is responsible for cloud formation and the absorption and emission of radiation by the atmosphere.

It has been recognized, that the ratio of the sensible heat flux to the latent heat flux, the Bowen ratio [248J

(180)

can be used to classify the underlying surface. Some values from different sources are compiled in Table 22.

The partitioning between the heat fluxes into the ground and into the atmosphere can roughly be estimated by the expression [255J

m! Im) _ ~l ku* 'l-'GR/'l-'SH- n-2--·

QCpku* y zow (181)


Table 22. The Bowen ratio ß for different surfaces

Surface ß Source

Desert dry 8 (20) 249 (250) semi-dry 3.5 249

Dry evergreen tropical forest, January 1970 6.4 251 Urban areas 2--4 249-251 Bare sand 1.0 249 Grass and farmland (cropland) in dry season 0.7 249-251 Savanna 0.6 249 Open land (arable land), moist 0.4 (0.2) 259 Dry evergreen tropical forest June 1970 0.45 251 Deciduous forests 0.33--D.4 252-249 Coniferous forests 0.5 249

wet canopy -0.3-+ 0.9 251 dry canopy

sunny -0.2-+ 10 251 overcast -0.3-+ 2 251

Sea -air interface AMTEX 1974 0.36 253

warm period 0.26 253 cold front passage 0.44 253

BOMEX 1969 0.09 254 Stable snowfield 0.1 251

Here (}s and (} are the densities of the soil or the ocean and the density of the atmosphere, cis the specific heat of either the soil or the ocean and cp the specific heat of the atmosphere at constant pressure. w is the circular frequency of the excitation, which is either determined by the diurnal or the annual wave; k and y are constant (k=OA, y= 1.781). u* is the wind velocity near the surface which is related to the friction height zoo Both quantities will be discussed in a later section.

The values for the ratio rpbrJrp~H lie between 0.5 and 1.9 for the diurnal cyc1e and between 0.1 and 0.04 for the annual cyc1e. High frequency variations generate fluxes which penetrate much more easily into the ground than into the atmosphere. The longer the period of the temperature variations at the surface is the better is the chance that the flux is directed into the atmosphere.

Heat Flux into the Ground

The heat transfer into the ground is a conductive flux which depends on the thermal properties of the material. The flux is proportional to the temperature gradient and the proportionality factor Je is the thermal conductivity measured in W m - 1 K - 1:

(182)

A flux propagating from a depth ZI to a depth Z2 is attenuated by the layer iflbetween which is heated by this flux divergence. The heating rate is proportional to the temperature change per unit interval multiplied by the mass contained in the layer and the specijic heat capacity c of the material.

224

It can be written

A4>GR=A.- -A.- =ca-,1z ( dT) (dT) iJT dz " dz " iJt

or for small differences

iJ4>GR = _ ~A. iJT)= -ca iJT. iJz &\ iJz iJt

If A. is taken as constant, the temperature change in the soil can be expressed by

iJT A. iJ2T iJ2T 7ft = Cii iJz2 =a iJz2'

where a=A./ea is the thermal dijJUsivity measured in m2 S-I.

H.-J. Bolle

(183)

(184)

(185)

Sometimes, e.g. under dear sky conditions, it is justified to apply the model of a periodical heating at the surface to the problem. The heating may be approximated by a sinoidal forcing functions with the circular frequency w:

T(O, t)= T+,17;,sinwt. (186)

T is the mean temperature, ,1 7;, the amplitude at the surface T(O, t). In this case the temperature in the depth z can be represented by

T(z, t)= T+,1T(z) sin[wt-lXz], (187)

where

,1 T(z) = ,1 7;, e-" amplitude at depth z, (188)

1/w . ffi· IX= V 2a attenuatlOn coe IClent. (189)

The attenuation eoeffieient IX, can be interpreted as the depth in which the amplitude is reduced to e- 1 ofthe initial value ,110, and may accordingly be called damping depth.

This solution iIIustrates some very remarkable features of the heat transport (see e.g. [256]). 1. The damping depth depends on the ratio of the exciting frequency and thermal diffusivity. Lower

frequency waves penetrate deeper than higher frequencies. A specific frequency penetrates deeper into a material of high thermal diffusivity (like stirred water) than into a thermal isolator (like organic material, peat).

2. Temperature ex trema occur if wt- V w/2a z = n/2. The time for a temperature wave to proceed from one level ZI to anotherone,z2, is accordingly: ,1t=(z2-z1)/V2wa.

3. From Eqs. (182) and (187) the expression

4>GR(O, t)= ,1 ToV meaA. sin(wt+"*) (190)

can be derived for the flux departing from the surface. The total flux into the surface can be determined from Eq. (190) if only that part of the phase is taken during which the sine is positive:

(191)

4. The heat flux at the surface is maximum for mt=n/4 [Eq. (190)]. The surface temperature (z=O) is maximum for mt = n/2. The maximum heat flux occurs one-eights of the period length prior to the surface temperature maximum. This is 3 h for the daily cyde and 1.5 months for the annual cyde.

5. For soils the damping or penetration depth for the diurnal wave is in the order of 20 cm (for dry sand) to 100 cm (for rocks). For the oceans it depends on the mixing of the surface layers. For the Gulf stream e.g. a depth of about 30 m has been estimated.

6. The annual wave penetrates 5-20 m into soils and about 600 m into the ocean of the type of the north-west Atlantic.


Flux of Sensible Heat into the Atmosphere

The transport ofheat into the atmosphere differs distinctly from the heat transport into the soil. As has already been mentioned in connection to the heat transport into the ocean, the thermal diffusivity becomes dependent on the mixing of the water, and for different oceanic states also different values are applicable for the thermal diffusivity a defined in the previous section. This is generally the case for liquids and gases. In an extremely thin layer on top of the soil or the ocean water, where no vertical wind component can develop, heat must still be transported by the molecules. But very soon a turbulent exchange process starts [257]. In tbis case the transport ofheat depends on the correlation of vertical wind speed fluctuations w' with the associated temperature fluctuations T':

(192)

Thequantity w'T' is difficult to measure and no theory exist to derive it directly from other parameters. The fluctuations responsible for turbulent exchange extend to frequencies higher than 1 Hz and are small in individual amplitudes so that elaborate techniques are necessary to determine the averaged cross correlations. In practise therefore often a formulation of the problem is applied which results from an analogy with heat conduction. The averaged heat flux from one level to another can be expected to be proportional to the temperature difference between these heights:

(193)

In order to account for temperature changes induced by adiabatic compression or expansion at vertical displacements, the pressure reduced potential temperature

e=Tc'~rcp, pinmb, (194)

should in principle be used far a discription of this process. e is the temperature which an air parcel with temperature T at pressure p will take if it is adiabatically compressed to 1,000 mb. This causes, however, problems if the layer is not "neutraBy" structured with an adiabatic temperature lapse rate. Since in practise always thin layers are considered for flux determinations, the potential temperature mayas weB be replaced by the normal temperature T. As proportionality factor in analogy to Eq. (182) the quantity (lepK is introduced as thermal conductivity to replace A, and x is an analogon to the termal diffusivity a. In the thin conductive layer just on top of the surface x == a a' the thermal molecular diffusivity of air, which takes values between 0.16 and 0.24 cm2 S-1. Above this layer where turbulent mixing processes take over, x has to be replaced by the eddy diffusivity K H for heat transport:

(195)

Since both processes, conduction and turbulent transport, exist simultanously in different magnitudes, x can in principle be regarded as the sum K H + aa. This,

226

700

600 City

500

E 400

-'=' '" 300 ..

I

200

100

0

7

5 0

N

-0 5 c

" ~ '" .. 4 > o .D

0 5

Pra ndl layer

U: Y..,.(ln) ..L k zo

~ 3 U:Uk• ;;;eu. -'=' ezo ----------'" ;; 2

10

Wind speed

Geostrophic w in d speed

10.4 ms- I

Suburban area

Wind speed (ms-I )

H.-J. Bolle

Rural area

Fig. 30. Vertical wind profiles over rough terrain after Hoffmann [258], and definition of roughnessparameter Zo

however, implies, that KH is a function of height rather than a constant. In fact an expression can be derived for K H from analogy to moment um transport, which makes use of the frequently observed logarithmic wind profile (Fig. 30):

KH~K=k2U(z)(ln z:r 1 ,Z>Zo. (196)

Zo is the roughness length which depends on the nature of the underlying surface. Hs numerical values vary between 0.001 cm for very flat surfaces, and e.g. 0.1- 15.4 cm for grass, 280 cm for a fir forest. u(z) is the horizontal wind speed


at height z which can be expressed by:

u(z)= u* ln~. k Zo

(197)

u* is the friction velocity which can be determined empirically from profile measurements. k is a constant, approximately 0.4 or elose to 1/e, the von Karman constant. For e-times Zo the wind velocity u(z)=e· u* (compare Fig. 30), and the speed u* occurs elose to Zo exp(1/e)= 1.44 zoo

There is reasoning about the validity of the assumption that the eddy diffusivity K H for heat can be determined from measurements of the momentum transport. For computation of the rate at which momentum is vertically transferred the following equivalent formulations can be used:

-,-, DU . 2 ,= -ew'u'=eK DZ =CrY2U . (198)

The first two relations are similar to Eqs. (192) and (193). The last term is mostly used for !lux calculations over water surfaces. CD is the drag coefficient [259J which generally is determined for a reference height of 10 m (over water: CD=2.6, 103 for UIO> 15 m/s; CD=0.5 .10- 3 . ~ for UIO< 15 m/s), compare also [260,261]. A new formula for thc drag coefficient over water has recently be derived by Amorocho and DeVries [262J which accounts for three different regimes: a constant slope region I for relatively quiet water surfaces (UIO < 7 m s- I), the transition region which starts with the on set of breakers and ranges to the saturation of breakers, and the constant slope region 11 for wind speeds (UlO) larger than 20 m S-1 The formula is

[ ( UIO -12.5)r' C,o=0.00151+exp - 1.56 ] +0,00104

and for other reference heights Z

[ 1/2 -I (1°)f 2 Cz = C,o I-(CIO ) k In --;]

Though differences have been found between K H and K under extreme situations there is generally no need for a distinction as long as the atmosphere remains elose to its neutral stratification (elose to the adiabatic temperature profile). From Eq. (195) a more simple formula can be derived which is often used in elimate studies.

(199) or

(200)

Here DH is the vertically integrated diffusivity a aa ~ K) -1 between the surface and

the height of the temperature measurement, e.g. 2 m, the standard meteorological observation height. In Eq. (200) use has been made of the relation (198)

( (ou/OZ)) Co=K ----:uz:- and L1u = u(z)-u(zo) = u(z).

More recent data are from Francey and Garratt [263] and Bill et al. [264] who applied eddy correlation method to flux measurements over sea and evaluated the bulk transfer coefficient Co for momentum, CH for sensible and CL for latent heat. Their results are presented in Table 23.

228

Table 23. Bulk transfer coefficients and fluxes measured over a water after Francey and Garratt [263] (a) and Bill et al. [264] (b)

Quantity

<PSH

<P LH

ß= <PSII/<PSH

CD CII

CL

Units (a)

(1.7±0.5)·1O- 3

(1.4±0.5) .10- 3

(1.8±0.3)·1O- 3

(b)

160±22 234±43

0.68 (1.75±0.26)·1O- 3

(1.17±O.l3) .10- 3

(1.19±O.l3) .10- 3

Flux of Latent Heat

H.-J. Bolle

Latent heat is essentially a transport of water vapor. If it condenses its heat of evaporation fi' is set free and transferred to the atmosphere as sensible heat. In order to evaporate water from the surface there must be water available and enough energy to initiate the phase change. This is always the ca se at the oceans or humid soils. In semi-arid soils efficient evaporation can only take place if the supply of water from deeper layers to the surface can compensate for the evaporation rate. The term "potential evaporation rate" is used if the water reservoir is large. The potential evaporation gives an upper limit for the evaporation and can be expressed as a function of air temperature, net radiation flux respectively sunshine dura ti on, water vapor partial pressure and wind speed [265,266].

The water vapor transport into the atmosphere can be described by formulas in analogy to Eqs. (192), (195), (199), and (200), e.g.:

<P LH = -Qfi'KL ~q ~0.622 Qfi'D (es-e), uZ p

(201)

where q is the specific humidity and e the water vapor partial pressure. The r.h.s. factor 0.622 is the ratio of the molar masses of water and air which results from the conversion of den si ti es to pressure. es is the saturation vapor pressure at the surface and e the vapor press ure in the atmosphere. The results of eddy flux measurements over water have recently been discussed by Anderson and Smith [267] who also compare a number of results obtained by different authors.

For c1imatological studies the flux of latent energy can be estimated over moist land surfaces and the oceans from the energy balance equation. If enough water is present the radiant energy is primarely used for evaporation. As long as the water supply is sufficient the surface will only slightly be warmed because the water has to be heated together with the soil material, and water conducts the heat immediately to deeper layers. Thus no strong temperature gradient between the surface and the adjacent air layers can build up, and the conductive transport of sensible heat into the lowest levels ofthe atmosphere, where heat can be picked up by the eddies, is very small. If therefore the sensible heat transport can be neglected together with the energy used for photosynthesis, Eq. (179) reduces to

(202)


Over land the flux into the soil is even smaller than the sensible heat flux so that here the term cJ>GR can also be neglected. In this case Eq. (197) reduces to

cJ>iB~cJ>~ (203)

and cJ>~ can be estimated via the Bowen ratio defined in Eq. (180):

oT JT _ cJ>~ _ QCrf(SH oZ '" cp Jz

ß- -1-- ---cJ>rn rOK oq - ro Jq

QoZ LH - oZ-oz Jz

(204)

if KSH=Krn=K, and ß~l. By successive approximation then cJ>rn can be recomputed from

cJ>iB = cJ>~ - cJ>~H . (205)

Over the oceans the latent heat flux is not exdusively determined by the radiation budget. It depends strongly on the advected heat in the oceans, and on the humidity gradient building up between the saturation pressure at the surface and the air. The energy for evaporation is partially drawn from the oceans themselves, and from the advected flux.

Energy Used for Photosynthesis

In chlorophyll containing plants the reaction

6C02 + 12H20 + hv -+ C6H 120 6 + 6H2 0 + 602 (206)

leads to the primary sugar products. Only rough estimates exist ofthe radiant energy consumed for photosynthesis.

This quantity can not be measured directly and its determination as residual among the other much larger components leads to inacceptable inaccuracies. The only reliable way to estimate the energy transferred to biological substance is via the consumption of CO2 . The fixation of 1 g CO2 in plants requires 10.5 MJjkg [268]. If Cis the transfer of CO2 mass in kg between the atmosphere and the plants, than

cJ>PH= 10.5· C MJ. (207)

From measurements of the CO2 concentrations Denmead [269] derived fluxes of CO2 mass into wheat and a pine forest. The corresponding short wave radiation fluxes are reproduced in Table 24. It can be seen from these numbers that 1-2.5% of the solar energy was consumed to build up biological substance. Baumgartner [249] arrives at an estimate of 0.13% efficiency of solar radiation for plant production on a global average. For the oceans the number is 0.07% and for land 0.3% (for individual stands as demonstrated by Table 24 up to 2.5%). The forests which cover 9% of the globe contribute 42% to the total biological mass production and consume 2.9 . 1021 Ja - 1 of solar energy. Budyko [270] estimates that 0.05% of the incomming total solar radiation or 0.1% of the photosynthetic active radiation is used for photosynthesis on the global average.

230

Table 24. Radiation components and photosynthesis

Plants

Wheat Pine forest

695 608

<1>*

378 461

Summary on tbe Partitioning of Energy at the Surface

8 15

0.023 0.025

H.-J. Bolle

The net radiant energy transferred to the earth surface is partitioned into three major fluxes, the turbulent sensible and latent heat fluxes into the atmosphere and the conductive heat flux into the soil or the conductive/turbulent transport into the ocean where eddy transports become important if the ocean waters are weIl mixed.

Sensible heat fluxes into the atmosphere get very large (about 93% of the radiation budget) only over dry deserts. Generally it remains in the order of 30% of <Pt over land and is in average still less over the oceans. During the winter months, however, it can locally exceed the radiation balance if heat is advected by warm currents underneath cold atmospheres thus building up a very steep temperature gradient between the surface and the air.

Maximum sensible heat fluxes are observed over African and Arabian deserts of 335 MJ m - 2 month - 1 in north summer as weIl as over the northern parts of the oceans over which very cold continental air masses or air masses originating from the ice sheets are advected. Here heat fluxes of 12.5 MJ m - 2 d - 1 occur. Values as high as 92 MJ cm - 2 day-l have been reported by Winston [271] in an outbreak of moist air over the Gulf of Alaska. Over the Gulf Stream 160-330 MJ m- 2

month - 1 and over Kuroshio 80-160 MJ m - 2 month - 1 and 2-20 MJ m - 2 d - 1

[253] have been determined. The latent heat flux is maximum over wet areas where it can reach 80% and

more of <Pt. It drops to about 7% over dry deserts. The latent heat flux is generally large over the oceans and can exceed the radiation budget by one order of magnitude over warm currents at high latitudes during the winter season if the heat is advected through the ocean.

The evaporation rates over the oceans are larger in winter than in summer because both the vapor pressure gradient and the wind speed tend to be larger during the cold season. Maximum values are observed over the Gulf Stream (920 MJ m - 2 month - 1 or in water mass to 11.4 kg m - 2 d - 1) and the Kuroshio (750 MJ m - 2 month -1 or up to 9.4 kg m - 2 d -1) in December. Minimum values of 40 MJ m - 2 month -1 occur at high latitudes in the North Atlantik ('" 70°) and North Pacific ('" 50°) during summer. Over land they are very small over the Sahara ( '" 0) and maximum over the east Amazones with '" 250 MJ m - 2 month - 1

during south winter. The evaporation is most steady and therefore high in an annual mean

(2.40 t m - 2 a - 1) in the trade wind zone where the wind is blowing with an average speed of 6-8 m s - 1.

The fluxes into the soil are generally small but not negligible. During the spring and summer months a small amount of heat is stored in the soil which is


released during autumn and winter. The storage in the oceans IS of larger amplitudes.

Energy Fluxes in the Atmosphere

Heat and Mechanical Energy Fluxes

Returning to Fig. 25 it is now necessary to consider the fluxes across the vertical boundaries of the atmospheric volume. Horizontal fluxes ot heat, kinetic and potential energy couple adjacent boxes and establish also teleconnections between far distant boxes in the respect that energy generated in one box may be transported over long distances and deposited in a far distant box.

The treatment of these advective fluxes goes back to Starr [272]. In energy exchange processes the following five species of energy playa röle:

qp = gnZ potential energy

v2 qK = - kinetic energy

2

qI=cvT internal energy (208)

qL = 2!q latent energy (q= specific humidity)

qw= E work done by press ure force against a vertical boundary. (]

The terms defined by Eq. (208) are per unit mass or specific energy species. The height Z which appears in the potential energy is measured in geopotential units. Because the acceleration due to gravity varies over the globe, the potential energy per unit mass is

qp= S g(ep, z) dz (209)

with z=geometrical height, cp=geographical latitude. With gn=9.80665 m S-2,

the standard acceleration of free fall, a new vertical scale, the geopotential height, can be defined by

1 Z = -Sg(ep, z) dz

gn (210)

so that (211)

The dependence of gon height can be neglected for the troposphere. gn is close to g (45°, 0)=9.8062m S-l.

Transports of the energy species defined by Eq. (208) are always accompanied by air mass transports or air mass exchanges. The air is the medium to advect the energy into regions where it is needed. The amount of energy transported depends therefore on the mass exchange which is measured in terms of wind velocity and air density.

The change L1Q in a reservoir Q of energy due to the transport of energy species defined per unit mass, qj, equals the displacement of the mass carrying this energy

232 H.-J. Bolle

times the specific energy. In case that the net mass change is zero:

L1Q=q*QL1V=q*QAL1s. (212)

A is the vertical area through which the energy is transported and L1s the horizontal displacement normal to A. Since the transport has to be considered as a mass exchange process, q* is the net value of the energy connected with inward and outward transports respectively:

q*= Ä (qi,in-qi,ouJ= Äqt=(Äqi)* q*=qin-qout. 1 1 1

The energy flux connected with this exchange is

n-. _ dQ _ * A ds - * A 'l"adv- dt -q Q dt -q QVn (213)

where Vn is the air velocity normal to A. If we consider the transport across the geographical latitude qJ through the

whole atmosphere, an incremental area dA can be expressed by

(214)

or, if the integration around the globe is performed with respect to dA, and dz is replaced by dp/gQ due to the hydrostatic condition, it folIows:

dA = 2nrE cos qJ dp . gQ

The total instantanous flux across the latitude is therefore

where

(215)

(216)

(217)

Here the terms cvTand p/Q have been combined to cpT, the sensible heat term, according to the First Law of Thermodynamics.

It is now of interest to consider energy transports averaged over certain time periods, e.g. a month, during which the directions of the transports undergo variations in space and time. Thus Eq. (216) has to be averaged over these periods. In order to establish such an average it is not possible to take average meteorological field parameters to compute the energy transports. Because of the exchange character of the transport it is necessary to compute first the instant transports for each incremental time interval and to sum them over the whole period. It can be shown that the zonal mean of the time averages can be represented by the sum of the following sub-averages: (1) Zonal average oftime averagedfield parameters. This term represents the mean

meridional circulation and would represent the energy transports if the atmospheric flow would be stationary.


(2) Zonal average of the deviations from mean field parameters. This term is =l=ü only, ifthere are perturbations which cross the area in different directions or at different speeds. This term represents travelling atmospheric systems.

(3) Zonal average ofthe deviation oftime averagedfield parametersfrom the zonal average. This term is =l=ü if there are zonal deviations of time averaged data from the zonal mean. This occurs if there are persistent regional perturbations of the field genera ted by standing eddies.

8

6

4

2

0

-2

§' -4 Potential energy a.. x -6 ::J

"" >- -8 Cl a;

-10 c: CI)

co 6 c: .2 Sensible heat :2 4 a; E c: 2 ca CI)

E 0 co

.--SE ----

::J c: -2 c:

<{

-4

-6

-1

0.1 -WE Kinetic energy

10 o 10 50 60 90 Degrees Latitude

Fig. 31. Energy transports in the northem hemisphere. (After Dort [274]). MMZ: mean meridional circulation (Hadley cell), SE: standing eddies (troughs, rigdes), WE: transient eddies (zyc1ones). Fullline: sum of fluxes due to MMZ, SE, and WE

Tab

le 2

5. E

xtre

me

valu

es o

fth

e en

ergy

tra

nspo

rts

in t

he a

tmos

pher

e o

fth

e no

rthe

rn h

emis

pher

e. (

Aft

er O

ort

[27

4])

Typ

e of

ene

rgy

Max

imum

of m

ean

annu

al t

rans

port

A

bsol

ute

max

imum

for

nor

thw

ard

tran

spor

ts

Lat

itud

e D

irec

tion

A

mou

nt(J

d-l )

L

atit

ude

Mon

th

Am

ount

(J d

-1

)

Pot

enti

al e

nerg

y 10

0N

N

6.

7.10

20

5°

N

Feb

r.

1.8.

102

1

Sens

ible

hea

t 10

° N

S

4.4

.102

0

Eq.

A

ug.

8.4.

102

0

45

°N

Jan.

5.

2.10

20

Lat

ent h

eat

35°

N

N

1.4.

102

0

Eq.

Ju

ly

3.9.

102

0

35°

N

Feb

r.

1.9.

102

0

Kin

etic

ene

rgy

30°

N

N

7.1

. 101

8

30°

N

Feb

r.

1.6.

101

9

Tot

al

45°

N

N

2.8.

102

0

50-5

5° N

D

ez.-

Jan.

4

.1.1

020

Abs

olut

e m

axim

um fo

r so

uthw

ard

tran

spor

ts

Lat

itud

e M

onth

A

mou

nt (J

d-

I )

Eq.

A

ug.

1.3

.102

1

50°

N

1.8

.102

0

5° N

F

ebr.

1

.2.1

021

5° N

F

ebr.

3

.9.1

020

Eq.

Ju

lyjA

ug.

2 .1

01

8

45°

N

May

1

.10

18

Eq.

A

ug.

8 .1

01

9

~

.j:>.

:I: ~ tIi

o ~


The energy transports can be determined from empirical data sets if there are enough informations on atmospheric temperature, humidity and wind profils available. This has so far only been accomplished with limited data sets [273-275J, especially with one data set gained during the International Geophysical Year (IGY) for the northern hemisphere.

The annual mean results of the study made by Oort are summarized in Fig. 31. The large transports of individual energy species partly compensate each other, especially in the tropics, and a total flux results with peak values of 2.8 . 1020 J d - 1

at 45° N. Table 25 provides some information on extremes which give also an indication of the reversal of the flux directions during the course of the year.

It can be seen that the mean meridional circulation contributes all of the potential energy and most of the sensible heat fluxes. Only latent heat and sensible heat at higher latitudes are to a larger degree also transported by eddies of which the transient eddies are the more important ones.

The kinetic energy fluxes are produced by pressure gradients which build up as a consequence of locally differential generation of potential energy. The kinetic energy therefore originates from potential energy. Kinetic energy is dissipated by friction. After Oort [274J 0.5 W m - 2 are dissipated in the mean meridional circulation and 1.8Wm- 2 in eddies. Wiin-Nielsen et al. [275J have estimated the annual variation of the total dissipation rate D (Table 26) which results in a slightly different annual mean value of2.1 W m- 2 .

Table 26. Dissipation rate for kinetic energy. (After Wiin-Nielsen et al. [275])

Month I 11 III IV V VI VII VIII IX X XI XII D in Wm - 2 2.6 2.4 1.9 1.5 1.4 1.5 1.3 1.5 2.3 2.5 3.0 3.4

If the value of 2.3 W m - 2 is accepted for the dissipation of kinetic energy which is 1.67--4.1.1020 W for the whole earth, then there must be a mean production rate of kinetic energy of the same magnitude. This energy must flow out of the reservoir of available energy. Consequently also potential and internal available energy must be produced with the same annual mean rate by the solar radiant input.

Oort [274J and other authors went even further into the analysis considering the distribution of the energy with respect to the mean meridional and eddy circulation systems according to the concept which has originally been developed by Lorenz [276].

In this concept the available energy is genera ted by a rate of 3.1 W m - 2 from radiant energy and is converted to mean meridional kinetic energy at a rate of Cz=O.1 W m- 2 as well as into eddy available energy at a rate of TA = 3.0 W m- 2.

The eddy available energy is converted into eddy kinetic energy at a rate of CE = 2.2 W m - 2 and dissipates at a rate GE = 0.8 W m - 2 due to temperature adjustment.

Between the kinetic energy of the eddies and the MMZ exist an exchange such that eddies supply the mean meridional circulation with kinetic energy at a rate of

236 H.-J. Bolle

Tz = 0.4 W m - 2. Finally the kinetic energy is dissipated at the rates of DE = 1.8 W m - 2 for the eddies and 0.5 W m - 2 for the MMZ as mentioned before.

It is more an academic question how large the total atmospheric energy content is, but it is of great interest how much of this total energy is available for energy transformations. If a reference atmosphere is considered which is in hydrostatic equilibrium, one could compute which amounts of energy would be released if the atmosphere is cooled to absolute zero and collapses to a thin layer. In that case approximately 1.17· 1024 J would flow out from potential energy and heat. Oort [274J has computed 5.875· 1023 J for the northern hemisphere. According to Oort [274J the variability of the energy content of the atmosphere is very smalI, about 2.3% compared with this total amount.

It is interesting to note that the amplitudes of these variations are in the following orders of magnitude for the individual energy components in the northern hemisphere:

Potential energy ±2 .1021 J

Internal energy ±6.5· 1021 J

Latent energy ±4 . 1021 J

Kinetic energy ± 1.2 . 1020 J

Potential, internal and latent energy reach their maximum value 1.5 months after solstice in summer with a minimum in January while kinetic energy has a maximum in winter and its minimum in summer.

Deposition of Energy in the Atmosphere

Radiative H eating and Cooling

If a radiative energy flux penetrates an atmospheric layer and interacts with the gases as weIl as the aerosols including clouds some energy may be transferred between the radiation flux and the air volume. Both effects occur: radiation is absorbed by the matter, and radiation emitted by the air volume is added to the flux. If the gas volume gains energy from the radiation field it is heated and if the emitted energy exceeds the absorbed energy the volume is cooling.

Quantitatively these relations can be expressed as folIows. A radiant flux density M (in W m - 2) penetrates for the time L1 t a surface A and enters a volume V=A·L1s.

After a distance L1s a flux density M + L1M emerges from the volume. Then the energy

L1Q = L1cP· L1t= L1M . A . L1t (218)

is exchanged between the radiation field and the volume. The change of energy within the volume can be expressed by the change in temperature L1 T and the specific heat (at constant press ure) of the air:

(219)


The mass m of the air in the volume A . Lls is

m=Q·A·Lls.

The three Eqs. (218)--(220) give

cpQLlsLlT=LlM. Llt.

The rate of the temperature change Ll TI Ll t is therefore

LlT Llt

Equation (222) can be generalized to

8T ot

1 LlM

_1- div S, QCp

where M == S is the Poynting vector of the radiation flux.

237

(220)

(221)

(222)

(223)

The heating respectively cooling rate expressed by Eq. (223) is a measure for the energy exchange between the atmosphere and the radiation field. Heating is generally induced by the absorption of solar radiation in the ultraviolett and near infrared part of the solar spectrum and by the aerosol absorption throughout the shortwave spectrum (see Fig. 11). Cooling on the other side is due to infrared emission at wavelengths beyond 3 !Jlll.

Vertical Distribution of Energ y Sour ces and Sinks

Surface

The earth's surface is an important regulator for the radiation processes in the boundary layer. Different situations exist over the ocean and the solid earth. The oceans absorb about 94% of the incident solar radiation and the outgoing infrared radiation remains very constant for aperiod in the order of several days. The ocean stores the heat in a rather large reservoir which changes its temperature only slowly. At the solid earth surface a smaller fraction of the solar energy is absorbed (~80%) and is not immediately transported into deeper layers. Therefore the surface heats up during daytime thus decreasing the radiation budget due to larger emission. During nighttime the surface cools because of the longwave emission. There exists a marked diurnal temperature wave, the amplitude of which depends on soil humidity and the fraction of water which is evaporated. Over snow and ice covered terrain or water the temperature wave is limited by the temperature of melting ice and snow so that in this case there is also a smaller temperature wave to be expected than for bare soil.

The surface is in any case a sink for solar radiation (even freshly fallen snow absorbs a few percent). For the longwave radiation, however, it is a source of radiant energy because of its high temperature compared with the atmosphere. The net effect of the surface depends on the temperature structure of the surface-atmosphere system and the magnitude of insolation. For annual mean conditions the earth surface is a sink for net radiant energy for latitudes at least up to about 70° and a source at higher latitudes (especially in the southern hemisphere).

Troposphere The chemical constituents of the troposphere are relatively weak absorbers in the visible part of the spectrum but some of them are strong absorbers in the ultraviolet and in the infrared. Ultraviolet radiation does, however, not penetrate the stratosphere in larger quantities. For solar radiation the tropophere is therefore a relatively small sink - the absorption concentrates mainly on the water vapor bands in the near infrared but also aerosols, c10uds and ozone contribute.

238 H.-I. Bolle

In the far infrared the situation is more complicated since the troposphere absorbs radiation emitted by the surface and emits radiation at lower temperatures. However, the irradiance which the atmosphere receives from the surface is to a large fraction compensated by its emission towards the surface. So the atmosphere still emits a substantial amount of energy to space. In fact this loss of radiant energy is much larger than the gain due to absorption. Therefore the troposphere is a net source for radiant energy.

Atmospheric gases are distributed according to the hydrostatic equation

(224)

where l!o is the density at altitude Zo and H is the scale height defined by Eq. (108). Water vapor is much more concentrated in the lower atmosphere than all the other gases. This circumstance can be accounted for by assigning to H20 a much smaller scale height (about 2 km) than for air (8 km). Also aerosols are concentrated in the atmospheric boundary layer with a scale height of about 1.5 km. The strongest losses of infrared energy due to gaseous and aerosol emission will therefore result from the lower troposphere. However, elouds change this picture since they are mostly placed on top of the bulk of water vapor and aerosols. They are furthermore the strongest atmospheric emitters. Only because the temperature decreases towards the tropopause and because high cirrus elouds are partly transparent their relative contribution to the emitted energy decreases slowly as the tropopause is approached.

The combined effect of direct radiant energy input into the atmosphere and its heating from the surface during daytime is used to warm the lowest layers of the atmosphere and to generate buoyancy, the result of which is turbulence and convection. Under the influence of this energy input the temperature profile of the atmosphere adjusts itself in the average to the adiabatic lapse rate of -6.5 K/1,OOO m for humid air. <

During night under elear sky conditions the lowest meters of the atmosphere are cooled which frequently results in a temperature inversion. If the humidity of the air is large this cooling may lead to condensation and fog accompanied by a strong cooling at the top of this layer. In areas of strong pollution the temperature inversions can be enhanced by the accumulation of aerosols at the top _of the boundary layer.

Also in the free atmosphere temperature inversions occur especially if the water vapor concentration suddenly drops or if eloud pre-condensation stages lead to higher emissivities. In these cases, where a radiating component sharply changes its concentration from high values at lower levels to low values, the atmosphere cooles in the region below the inversion. Cooling rates of a few degrees are computed. If there are elouds in the atmosphere then the main cooling surface is elevated to the top of the cIouds where maximum infrared cooling occurs. On the other hand also most of the solar energy is deposited during daytime in the upper layers of the cloud. While - in a study of Stephens [89] - the longwave cooling of a Sc layer reaches about 5 K at the top of the eloud, the heating by solar radiation in the same layers is only about 2 K but extends much furtner down towards the bottom of the cloud. The net effect is therefore a cooling of about 4 K at the top of the eloud and a heating of about 0.5 K at its bottom which is partly also due to the absorption of infrared radiation from the ground. The situation varies from eloud type to cloud type and no general statement can yet be made. The investigation of radiative properties and effects of clouds is presently a major research project within the World Climate Research Programme.

With the condensation of water vapor latent heat is deposited in the eloud levels between about 2 and 7 km, higher up in the tropics than at high latitudes. The heating rates attributed to latent heat release are maximum at the equator at about 6 km altitude and reach here nearly 3 K in a zonal average. The effect of this heating by latent heat release extends to the poles [273J.

Sensible heat is transported 1-2 km high into the boundary layer where it contributes to the heating rate from the equator to latitudes around 70°. This contribution is maximum in the latitude zones 30-40° N and around 30° S where it is up to 2 K day-l on a zonal average elose to the surface. The heating over the oceans due to sensible heat fluxes is much less than over land.

The röle of aerosols is discussed in more detaillater.

Stratosphere Above the tropopause the ozone concentration starts to increase until a maximum is reached around 25 km. The emission of the ozone between 9 and 10 J.UIl wavelength is a prominent feature in the


terrestriallong wave speetrum (see Figs. 15 and 16). The effeet of this emission is a eooling of the layers 30--65 km and a slight warming of the tropical and midlatitude atmospheres between the surfaee and 30 km (heating maximum around 18 km).

Carbon dioxide - beeause of its eonstant mixing ratio - eools throughout the stratosphere (and troposphere) due to its emission at 15 J..Ull and 4.3 J..Ull. The water vapor eoneentration drops sharply at the tropopause but still eontributes subtantially to radiative eooling beeause of its extended infrared speetrum. In the lower stratosphere there exist therefore a strong eooling effeet. The net eoöling dueto long wave emission in the 74 mb (~18 km) layer is mueh stronger (~7 K) at higher latitudes (70°) than in the tropieal zone ( ~ 2.5 K). In this layer ozone is still heating.

The energy losses of the stratosphere due to long wave radiation are almost eompletely eompensated by daytime heating due to shortwave absorption of solar radiation in the ozone bands. COz and HzO both do not have absorption bands in the speetral region where solar irradianee is maximum and the water vapor eoneentration above the tropopause is so low that its near infrared bands do not any more eontribute substantially (e.g. [77]) to the radiation budget. X-rays between 10- 3 and 7 .1O- z nm eontribute between 27 and 50 km to heating by absorption in Ar, 0, N, and C.

In eombination of all these effeets there results a small eooling of 0.5-1 K in the winter hemisphere while in the summer hemisphere the radiation budget is almost in balance between 14 and 25 km altitude [277,273].

Mesosphere In the mesosphere again X-ray radiation between 0.07---0.7 nm is absorbed by all atoms. In addition absorption oeeurs in the Sehumann-Runge dissociation eontinuum of oxygen (135-175 nm) of whieh ~ 30% is eonverted to heat. Also 0 3 starts here to absorb wavelengths > 100 nm. The eooling by ozone and COz eontinues throughout the mesosphere with COz as the dominating gas. The ozone heating rates beeome almost zero above 65 km. The eontribution of HzO is not weil established. There may still be some eontribution to eooling in the lower mesosphere. Near the thermopause ( ~ 80 km) the eooling as weil as the heating due to ozone absorption are strongly redueed: the ozone absorption beeause of photodissoziation of 0 3, and the COz eooling beeause of the temperature strueture in this region. Around 80 km there might even be a layer of about 5 km thiekness in which heating due to infrared COz radiation oeeurs. Beeause of its strong dependenee on solar radiation heating is redueed in the winter hemisphere. Over the winter pole there is a net eooling effeet of about -12 K of 65 km altitude. The summer pole mesosphere is heated by a rate of 6 K at the same altitudes and the warming is extending to the tropies of the winter hemisphere. The resulting strong gradient gives rise to latitudinal eireulations.

Thermosphere In the upper thermosphere solar energy is absorbed during daytime by atomie oxygen and nitrogen due to photoionisation by which an eleetron gas of a broad energy speetrum (0 to ~ 70 eV) is genera ted. The maximum e1eetron gas heating rate' oeeurs around 175 km altitude. In the layers around 225 km 0+(4S) is generated by radiation of ,1,<91 nm, whieh reaets with N z and Oz to produee NO+, Ni, and Oi ions.

In the lower thermosphere radiation ofwavelengths between 27.5 and 79.6 nm direetly produces ions like Ni, Oi, 0+, and NO+ due to absorption by the neutral atoms respeetively moleeules. Around 140 km radiation between 15 and 24.5 nm is absorbed by N z, Oz, and 0, as weil as in the Lyman eontinuum (79.6-91 nm) of 0 and Oz. Furthermore Oz is ionized by radiation between 91.1-102.8 nm including the strong Ly ß (102.5 nm) and qIII) (97.7 nm) lines of the solar speetrum. Also X-ray radiation of 3.5-8 nm produces ions here. The result is a rather strong heating rate around 100 km ofabout 10- 13 J em - 3 S-l or 9 Kjday, to whieh also Joule heating due to ionospherie eurrents eontributes - especially in the ease of Aurora.

The total possible heating of the thermosphere depends on the value of the available very shortwave solar e1eetromagnetie energy, whieh ean vary by a faetor of 2 or 3. The aetual energy input to the thermosphere depends on the solar zenith angle. This gives rise to strong temperature variations and eonsequently mass transports between the day and the night hemispheres, as weil as to a superimposed large seale transport between summer and winter hemisphere.

240 H.-J. Bolle

The heat input to the upper atmosphere has to be dissipated somehow. In the reaction chain involved there are generally secondary ionisations and exitations by the electron gas, and emission of excited molecules and atoms. As has been pointed out by Thomas [278] in a review on this matter the excited molecule N! seems to play an important röle in this context since it cannot radiate directly and is not readily quenched by collisions with other molecules so that the transfer into translational energy by collisions with 0 and the excitation of CO2 or an ion-atom interchange to produce NO may be important for the energy budget of the lower thermosphere.

The consideration of infrared emission by CO2 becomes very complicated in the lower thermosphere because firstly not much is known about the variations of CO2 concentrations in this region, and secondly the emission is of the type of non - thermodynamical equilibrium [280]. Cooling by CO2 emission occurs with a peak at about 90 km and vanishes above 120 km because of CO2

dissociation. At higher altitudes the atomic oxygen emission at 63 J.UIl produces a growing contribution to the cooling [281].

It turns out, however, that the energy budget of the thermosphere cannot be maintained by these emissions alone, and additional dissipation must occur. One possibility is the downward heat transport by turbulence into the mesosphere, or additional radiation to space as has been suggested by Markov [279] in the 3-8 J.UIl region. At about 100 km also the recombination of 0 to O2 contributes to the budget of the thermosphere as well as the dissipation of waves (Mayr et al. [282]).

Energy Transports and Exchanges it the Atmosphere-Ocean System

General Remarks

The response of the earth-atmosphere system to the energy input cari be understood as an interplay between a number of processes on different scales by which energy is converted and transported in order to establish a balance between the external forcing provided by the astronomical configuration and the internal dynamies. On earth exist two large scale transport systems: the general circulation of the atmosphere and that of the oceans. Both serve adjustments between energy reservoirs. To maintain the c1imate of the earth it is necessary that the budgets of the transports of energy, mass and momentum are balanced within a certain period. This period is determined by the geometrical configuration between sun and earth, which governs the cyc1e of the energy input, and is thus one year or, if one allows for small interannual variations, a few years.

Since the orientation of the spin axis of the earth is nearly vertical to the ec1iptic, the annual mean input per unit area of solar energy is maximum in the tropics and minimum at the poles. On the other hand the earth looses energy by emission of long-wave infrared radiation at alliatitudes. In the polar regions the annual loss is larger than the annual solar radiant exposure, while in the tropics more solar radiation is gained during the year than is emitted in the infrared.

A local balance of the radiation fluxes would require higher effective emission temperatures in the tropics and lower mean effective emission temperatures at the poles. One reason that such a temperature regime is not feasible is the actually strong seasonal dependence of insolation at the poles, because the earth axis is not exactly perpendicular to its orbital plane. During the summer season the extraterrestrial solar radiant exposure of the polar regions is even slightly larger than at the equator.


Instead ofbeing locally balanced the atmosphere and the oceans transport the required energy as heat away from the equatorial regions for most of the year. This heat transport requires strong mass transports respectively mass exchanges which are accompanied by momentum transports and transfers between solid earth respectively oceans and the atmosphere. These transports are partly responsible for the regional climates and for the atmospheric circulation systems, which develop in response to this seasonally variable external forcing.

Circulation Pattern in the Atmosphere and in the Oceans

The air heated by the strong energy input in the tropics rises and splits into one northward bounded and one southward bounded branch. During the upward motion the air masses are dryed out since the water vapor condenses in large cummulonimubus clouds that cluster in the "Inter-Tropical Convergence Zone" (ITCZ) which can easily be located in satellite images as a band of high reaching bright clouds near the equator. In the northern summer the ITCZ is found at 10-15° N and during northern winter around 5° N. That the ITCZ does not have its mean position at the equator is due to the different land-sea distribution in the two hemispheres.

The rising air gradually cools, the heat being converted into potential energy. Near the surface in the planetary boundary layer, warm and wet air is entrained into this vertical movement so that here in the lower troposphere an equatorward flow developes. This flow is deflected into a wind system with a westward component due to the Coriolis force generated by the rotation of the earth. This force acts always at right angle to the direction of motion, towards the right in the northern hemisphere and towards the left in the southern hemisphere. The regions in which this south east respectively north east winds develop are called the trade wind regions. They are disconnected from the flow in the upper troposphere by a temperature inversion of a few hundred meters to 1-2 km height.

Thus far we have seen that the air from higher latitudes which flows towards the equator in the boundary layer is warmed up due to sensible heat input from the surface and latent heat release in large cloud towers. It rises at the "meteorological equator", the ITCZ. We now have to look for the closure of this circulation. At a non rotating planet the air would sink at the coldest area which is the anti-solar point. If a planet rotates very slowly like Venus (rotation period 1/250 days) the descending branch ofthe air will be near the poles. The reason for the break-up of the large circulation "celI" on earth lies in the action of the more vigorous Coriolis force. The dynamics of a steady gas flow on a rotating planet requires that the Coriolis force is balanced by apressure gradient force (and a frictional force near the surface). To maintain a flow from the tropics to the poles would e.g. require that the pressure grows always towards the east around the earth in the northern hemisphere, which is impossible. Such a flow would also be in contradicion with the momentum budget. Instead, the motion in the free troposphere is bent into an eastward directed wind system around 30° latitude which ceases the direct transport towards the poles. The now cold air of the upper troposphere is sinking in this zone and a fraction of it returns directly at lower

242 H.-J. Bolle

levels to the tropics. The large scale circulation which has so far been described is called the Hadley [283J circulation, which is the backbone of the earlier mentioned mean meridian circulation [284]. The Hadley cell is, however, not a permanent but only a mean feature of the atmospheric circulation. The instantenous structure of the circulation is much more complicated with sinking motions inbetween the cumulonimbus cells and sometimes extended rising motions up to 200 N [285].

At the high latitude boundaries of the Hadley ceIls, the polar fronts, a completely different regime takes over the energy transport. Here systems with vertical rotation axes develop: high and low pressure systems which are partly stationary and partly quickly moving.

The stationary systems group in a wave like pattern which moves slowly around the earth. The reason for these waves is, that a zonal air flow carries a certain angular velocity which results from the rotation of the earth and from the curvature of the flow as weIl. This "vorticity" is a conservative quantity. The Coriolis force, which is responsible for the curvature of the flow due to the earth rotation, is increasing towards the poles with the sine of the geographical latitude. If an air flow from lower latitudes is directed into higher latitudes the Coriolis term is increased at the expense of the curvature of the flow, since the absolute angular momentum of the air parcel, the absolute vorticity, has to be conserved. If this curvature was zero at one latitude and the air parcel is moved polward, then the flow has to adopt a curvature with negative vorticity, against the direction of earth rotation, or "anticyclonic" (like in a high pressure system). As this .(negative) curvature increases with the polward motion the flow will finally bend back towards the equator, the Coriolis term is reduced as is the negative curvature. The flow finally reaches a point at a lower latitude where its relative vorticity gets positiv the curvature changes its sign and gets cyclonic which reverses the process.

Warm tropical air is flowing in with this wave whenever the flow is directed polwards where high press ure "ridges" build up and cold polar air is carried back in the equatorwards directed "troughs" of low pressure: slowly moving or even standing eddies of large extent. This exchange of cold high latitude and warm tropical air in these latitudes is a much more efficient energy transport mechanisms than a cell with a horizontal axis could be. The flow pattern is called Rossby waves. The Rossby-wave zone is therefore a region where cold and warm air masses merge so that large temperature gradients develop and a ''front'' builds up. At this polar front the quickly moving transient eddys are generated: the cyclones which are a dominating feature of the weather at mid latitudes.

In the upper troposphere air from the high lying tropical tropopause, the boarder between troposphere and stratosphere, warms in its slowly sinking motion andconfluents withrelativelycoldairfrom the polar region, and strong temperature gradients develop. In these beIts the subtropical and polar-front jet streams deveIop: strong winds of about 100 m S-l velocity which are generated by the strong thermal contrasts. The subtropical jet is a rather frequent phenomenon between 30 and 400 latitude in the 200 mb level (12 km altitude). The polar front jet is highly variable, lies lower than the subtropical jet ( '" 9 km) and at slightly higher latitudes. It is connected with the polar front and the generation of dis turban ces.


Very low temperatures contrast at the tropopause level in the tropics the relatively warm temperatures at the same altitude in the polar lower stratosphere, especially during summer. As one moves from the troposphere upward into the lower stratosphere therefore the temperature gradient changes direction in the polar front region. This causes the jet streams to cease in the lower stratosphere. Also the precise tropopause level is undetermined in this region, which is called the tropopause break.

The circulation over the polar caps consists of cells with cyclonic (counterclockwise in the northern hemisphere) and anticyclonic (clockwise in the northern hemisphere, in the southern hemisphere reversed) circulations. These cells, which are piloted by the land-sea distribution, provide for the efficient exchange of polar and non-polar air masses.

This large scale picture of the general atmospheric circulation system is further complicated at different scales by wind systems which develop regionally and seasonally because of differential heating. The most prominent circulation systems at the scale of continents are the monsoons. During summer the continents and especially highlands like Tibet warm up generating rising air (low pressure at the surface). The outflow of this air at higher levels is sometimes injected into large scale zonal transports before sinking down over relative cold parts of the oceans. The air over the hot areas of the continents have to be replaced by air from the adjacent comparatively cooler oceans. This air however, is nearly saturated with moisture and if it starts to raise over mountains very intense precipitation develops.

In winter the situation is reversed: the continents cool off rapidly because of their low heat storage capacity and the air on top of them get into a sinking motion flowing out into regions over the oceans. The oceans are still of approximately the same temperature as during summer time but are now warmer than the continents so that the air raises over the oceans.

Circulation of the same type but at much smaller scales develop daily over shoreIines generating a sea-Iand circulation known as sea breeze at daytime. In mountain areas small scale circulations develop in valleys where the slopes exposed to solar radiation warm up much more rapidly than northward directed slopes and the bottom of the valley. At nightime cold air from the higher layers sinks down to the bottom of the valley and produce "cold air lakes".

It was believed that the oceans participate in the large scale energy transport in a similar manner as the atmosphere. The heat is mainly deposited in the tropics. New satellite measurements indicate, in fact, that the heat input in the tropics may even be larger then previously estimated from conventional meteorological observations: the cloud coverage seems to be less than these earlier estimates showed. The heat genera ted by the solar irradiance is deposited in a layer of some decameters to one - to two hund red meters thickness down to a layer where a steep decrease in temperature in the order of several degrees occurs, the socalled thermocline. This phenomenon can be looked at as analogous to the atmospheric inversion on top of the trade wind zone.

In the upper layer of the ocean currents are genera ted by the action of the wind on the surface. As has first been derived theoretically by Ekman [2J the ocean transports generated by the wind forces are not parallel to the wind but are directed in the northern hemisphere to the right of the wind direction. This is

244 H.-J. Bolle

again a result of the Coriolis force. At its surface the ocean flow will be under an angle of theoretically 45° to the wind direction (in practice the angle is smaller, in the order of 15°) and the flow will veer with depth to 180° until the effect of the wind vanishes because of internal friction. The Coriolis force acts at 90° off the wind direction on the center of gravity of the moved masses. Therefore the effective transport is at right angle to the wind vector for this center of gravity, to the right in the northern hemisphere and to the left in the southern hemisphere.

This has interesting consequences at the ITCZ where the trade winds converge. The trade wind generate generally a current directed to the north-west north ofthe equator and to the south-west south of the equator. This causes a divergence in the ocean at the equator with the result that waters from deeper layer have to weIl up in order to replace the water masses at the surface. Cold water will then appear at the equator. If the ITCZ is north of the equator the wind from the south will cross the equator and then generate an ocean current towards the north east. This causes a much more complicated structure in the currents near the equator. The ocean circulation system near the equator is shown in Fig. 32. The action of the wind on the ocean surface is referred to as Ekman pumping if it causes an upwelling of oceans waters.

At higher latitudes where we have a west wind regime the frictional force of the wind on the ocean surface is directed eastward and ocean currents directed towards the equator are generated. These will confluent with the polwards directed water masses of the trade wind zone in the polar front area. This convergence results in a downward and eastward directed motion at these latitudes.

Ocean currents can not develop freely like atmospheric wind systems, since the oceans are surrounded by the continents. By the action of the winds the surface waters are blown off the west coasts of the continents in the 0-25° belts. This gives rise to upwelling in these regions because the water masses have to be replaced from deeper layers.

At the east co asts, however, wind driven ocean surface currents are directed into a coast parallel flow in the 0-250 belt. Here warm currents develop which can eventually extend into mid latitudes. Such a situation is given for the strong Gulf Stream in the Atlantic and the Kuroshio in the West Pacific. These currents transport most of the heat from the equatorial regions into mid latitudes.

Ocean currents are also indirectly affected by the structure of the land near the coasts. In South America the high Cordilliers affect the atmospheric circulation system by deflecting the winds into a coast parallel flow. Instead of a south east wind as can be expected for the undisturbed flow, a south wind develops, which moves the surface waters right away from the coast (to the left of the air flow). This generates exceptionally strong surface currents into the mid Pacific and compensating upwelling at the coast. In this region (off the coast of Peru and North Chile) therefore low ocean temperatures and high temperature gradients prevail between ocean surface and atmosphere. Every year around Christmas ("EI Nifio"), however, warm tropical surface water moves into the upwelling area off the coast of Peru. In intervals of two to ten or more years this invasion of nu trient poor warm water intensifies considerably. It then remains in that area for most of the year - sometimes even for more than one year - and disrupts the economically very important anchovy fishery near the coast. As far as

Po

lor

fro

nt

150

E I

~O

u :c 13

0 0

1 '" .c. '" u C

l ~

::>

(/)

No

rth

eq

uo

tor

iol

rid

ge

0,

E---

E---E

---W

---W

-E--

"I/ .

....

... ,.W

·

200

, \ I

, ,

E

\ I

\ E

I I

\ I '

I,

, \

Eq

uo

tori

ol

rid

ge

S

ou

th e

qu

oto

rio

l ri

dg

e

Pa

cif

ic

SE

C

Pol

or

fro

nt

-W-

--W

--W

---W

---E

--

E-

-E

••

• ::.

. -

--

-)

, ,

\ I//"

\ \

-..

.•. .:

.....

....

..

-....

I

I \

I

--W

---_

W'"

'\

,I

;

I •

I,

......

" \

'00

I-\

\ ,I

\

\

" / ...

.......

..... .

....

...,.

" \

I ,,/

""

I

I \.

60

0 .....

.........

~ "

" ....

Atl

an

tic

m

BO

O 60o

N

o L

ati

lud

e

Fig

.32.

Mer

idio

nal

cros

s se

ctio

n o

f th

e m

ean

Atl

anti

c ci

rcul

atio

n sy

stem

aft

er D

ietr

ich

et a

l. [2

86]

(low

er p

art)

com

bine

d w

ith

the

zona

lly

aver

aged

dyn

amic

i"

~ ~. g [ tI1 iil '"' ~ '"" '"' '" ::; .g ::l- 5' ~ tI1 ~ g o ~ [ Vl

'< ~

heig

ht p

rofi

le (

rela

tive

to

500

dk)

of

the

Paci

fic

Oce

an a

fter

Wyr

tki

[287

b,

f, 28

9]

~

v.

246 H.-J. Bolle

recorded this happened 1891, 1925, 1941, 1957/58, 1965, 1972/73, and 1976. The phenomenon seems to be part of an almost global disturbation of the general circulation with strong zonal components in the tropics, such as have been realized the first time by Walker (compare [287 c-f]). The North Pacific trade winds strengthen while the South Pacific trade winds become weaker thus displacing the ITCZ slightly southward, closer to the equator than usual. Heavy rainfall occurs with this shift of the meteorological equator in equatorial areas which normally have a sparse precipitation record. Sea surface press ure drops in the eastern Pacific and the sea level rises. Due to changes in the cloud distribution and in the temperature gradients at the surface and the wind field, the energy exchange between the ocean and the atmosphere is strongly affected. The abnormal high sea surface temperature and associated phenomena spread out westward over the Pacific until it fades out in less than two years ("Southern Oscillation"). The reasons for and consequences of this anomaly are not yet weIl understood. There seem to be interrelations with climate excursions like strong winter events and droughts in different parts of the world, and even variations of the polar energy budgets may playa röle.

Heat Transport by the Oceans

There are only few direct observations of heat transports in the oceans available. This would require an observation system for ocean temperatures and flow velocities similar to the aerological observation system for the atmosphere. Estimates of the heat transport have therefore partly to rely on indirect methods.

One method is to determine the oceanic heat flux as residium of the total flux determined from satellites if the transport of the atmosphere is known. The first attempt to apply this method was given in Fig. 31. The method was later on refined and resulted in the values of Table 27 for the northern hemisphere.

Table 27. Heat transport by the oceans after Oort and Vonder Haar [214]. 1 PW = 1015 W

Latitude

Heat flux in the oceans across latitude circle in PW residual method

0.20 1.05 2.90 2.00 1.90 0.90 -0.20

Another approach is the evaluation of all available data on heat exchange between the oceans and the atmosphere. At one latitude the heat flux has to be assumed zero and from this latitude on the input of solar radiation and the other energy transfers at the ocean - atmosphere interface as weIl as the divergence of the total heat input has to be computed for each latitude. The values derived in this way by Hastenrath [288], reproduced in Fig. 33, compare favourable with the Oort and Vonder Haar da ta mentioned before.

Bryden and Hall [289] determined a value of 1.11 PW at 24° N in the Atlantic for which latitude Hastenrath's value is 1.55 PW. The evaluation of direct oceanographic measurements led Roemmich [290] recently to the value of 1.2 PW


66 _~ --66-~0_ ( ________________ ~ r~/N~I/i.. 60N~-----------------~~r~

v

20 E

, I 1 1

I~ 77ma.~

49 -~---

: 17 J'

120 E o

Fig. 33. Annual mean oceanic heat transport in units of 1013 W adopted from Hastenrath (288)

247

for 24° N and 0.8 PW at 36° N and a nonzero northward flow at 48 N the magnitude of which could not be determined with a sufficient degree of accuracy (in the order of 0.5 PW).

The uncertainties involved in all of these methods is still larger then can be accepted for climate studies. Nevertheless the application of these methods and a careful analysis of their results led to a revised picture of the heat transport in the oceans [288, 291]. Apparently the Pacific ocean is the only one which behaves "normal" in the sense that it transports heat symmetrically out of the equatorial region into both hemispheres. The Indian ocean is blocked in the north at about 20° latitude by the Indian continent. Most of the energy gained in the tropical zone therefore flows out in southward direction and is tran spor ted into the Atlantic ocean via the Anguellas current. Also the Pacific joins with the Indian ocean in supporting the Atlantic Ocean's heat budget. The tropical Atlantic therefore exports heat mainly to the northern hemisphere and there exists a strong transport across the equator in the Atlantic ocean. This picture is still a working hypo thesis which has to be verified by more direct methods.

As has been demonstrated by Washington et al. [292], the oceanic transport can as weIl be splitted into a transport by mean meridional circulations, eddies and diffusive transport (Fig. 34).

Energy Budget of the Earth-Atmosphere System

The total solar power intersecting with the earth is 174 PW. 9 Of this amount 69-70% are absorbed and used for energy conversion processes on earth. A little more than one third of these 69% is absorbed in the atmosphere. The

9 1 pw= 1 Petawatt= 1015 W

248

3: Il..

c

x :>

0 <I>

:r:

3

2

0

-1

-2

-3

/ l i i

;i -- Meridional Transport !

Gyre Transport ....... \ ...... \.'-.J;;';! -4 Diffusion Transport

Total Transport

-5.~~~-~-~-~-~~_\~'··L.!_;··_·:~~~ 80 N 60 S 80

Latitude

Fig.34. Components of the heat transport in the oceans. (After Washington [292])

H.-J. Bolle

Table 28. Estimated relative importance of different atmospheric constituents for the earth's global mean radiation budget

Shortwave radiation (0.3~ 3 filll)

Solar extraterrestrial flux at 1 A.U.

Reflected sunlight Aerosols Surface Gases Clouds

Absorption of solar radiation

Surface Open oceans Open land Iceand snow

Gases H 20 0 3,02

CO2

Others Clouds Aerosols

1% 4--5% 5~6%

19%

30-31% 9%

4~ 5%

12% 5~ 6% 1%

< 1%

100% = 174.6 PW

30-31 %

69~70%

43-45%

18~20%

4~ 5% 3~ 5%

Longwave radiation (3~ 100 filll)

Emitted IR flux Clouds Gases

H 20 CO2

0 3, CH4

Surface Aerosols

16 7

100% 47~50%

26~27%

18% 8% 2%

23~24%

1~ 2%


major part penetrates to the surface and is absorbed in the oceans and by the land surface. The approximate fractions are given in Table 28.

From these figures one can es ti mate the heat budget of the atmosphere. 25% of the extraterrestrial solar radiation or 43.5 PW is absorbed in the atmosphere, and on the other hand about 77% of the infrared emission which is equivalent to 54% of the extraterrestrial solar radiation results from the atmosphere. The loss of radiative energy from the atmosphere is therefore

(0.25--0.54) ·174.6= -51 PW.

Applying Eq. (223) the cooling rate of the whole atmosphere can be estimated if for the volume of the atmosphere the earth surface (1.27 . 1014 m2) times the scale height (8,000 m) according to Eq. (108) is used. With an air density of 1.275 kg m -3

and a specific heat capacity of 1 J K -1 kg- 1 a cooling rate of little more than 0.8 K d - 1 is obtained. More accurate estimates which take e.g. into account the volume lost by mountains, result in a cooling rate of about 0.9 K d - 1.

In order to keep the atmosphere in thermal balance the fluxes of sensible and latent heat have to compensate for this loss (Fig. 35).

The fluxes of sensible heat are smaller over the oceans than over the continents. A rough estimate from Budyko's maps yields an average flux from the oceans in the order of 4 . 108 J m -2 a - 1 and over land of 10 . 108 J m -2 a - 1.

F or the latent heat the corresponding numbers are 30 and 10· 108 J m - 2 a - 1

respectively. The average Bowen ratio for the oceans is therefore in the order of 0.1 and for the earth surface in the order of 1. More accurate estimates for the oceans yield aBowen ratio wh ich depends on latitude from about 0.1 in the tropics to 0.45 at 70° north and 0.23 at 70° S [293].

The latitudinal distribution of the heat fluxes for the northern hemisphere is schematically drawn in Fig. 36 using the data of Oort [274]. All data are annual mean va lues. The uppermost area represents the outgoing infra red flux per 5°

100

30.5

Fig. 35. Energy budget of the atmosphere. Numbers are in percent of the solar irradiance at the top of the atmosphere

250

VI CL> X :::l

>-01 L-

CL> e CL>

Cl :::l e e

<{

(oS/O/<l1 nOr zzol ·SSl ·\: =)

auoz ·101 oS /1l0Ms10l

"Oft! § !I !I

H.-J. Bolle

CL> ~ .c: -'" e -0

- e L-

0 VI a. e VI Cl e Cl CL> L- U f- 0

~

Fig. 36. Annual zonally averaged energy fluxes in the earth-atmosphere system of the northern hemisphere. (After numbers given by Oort [274])

0

'"

0 <Xl

0 .....

0 1.0

0

'"


latitude zone in PW, the seeond area from top the distribution of absorbed solar radiation. The two areas are of equal size. At about 30° N the point markes the latitude where both amounts equal eaeh other. South of this erossover point the absorbed solar radiation is larger than the emitted infrared radiation and viee versa. If the absorbed solar radiation is taken as 100% (= 69-70% of total solar energy reaehing the earth), then about one third of this amount is direetly transferred to the atmosphere and two thirds to the surfaee. Of these latter two thirds one part is direetly emitted to spaee from the surfaee of the earth. The net emission (earth surfaee emission aYs4 minus atmospheric infrared radiation ([>iw,o) amounts to about 24% (related to absorbed solar radiation = 100%). Another part of the energy is direetly eonverted to heat at the surfaee and released in situ into the atmosphere where it is transported by the mean meridional eireulation, standing and transient eddies whieh are symbolized by stream funetions and weather maps in the middle seetion of the figure. These transports are represented by the third area from the top in Fig. 36.

The water vapor transports have in great detail be investigated by Rakipova [294] and more reeently by Peixoto et al. [295]. Their zonally averaged results for meridional transports are presented in Fig. 37.

20 /'\

/ \ I \ I ,

15 : , I \ I ,.,-\ \ I! '. \ I! \,

10 I! \ \ li \ \ I; .. --- \ \ \

' VI Ii~' \ \ 5 - I; \ ,

'jf :/. \ , \ \ I \ \ I

'" o

2

~ ,~ x O~-----+~----~r-------------------------~~~~ 0 :>

o 0-e >

e 3:

S \ \

10 - \ \.

IS

I S IO

'\ \

I 0

-- 60 month s _ .- Wln er

_ ... - Spring

--- Summer •• -.---- Autumn

2

80 90 N

3: u.

x :>

ä CI>

.c:

~ CI>

ä -'

Fig. 37. Zonally averaged water vapor respective1y latend heat transport in the northern hemisphere. (After Peixoto et al. [295]). The 60 months period is [rom 1. May 1958 to 30. April 1963

252 H.-J. Bolle

The rest of the solar radiation is first absorbed in the oceans, transported northward and then released into the atmosphere (bottom area in Fig. 36), mainly as latent heat.

The radiation budget at the top ofthe atmosphere reflects the different surface conditions. The seasonal variation of the radiation budget is primarely determined by the solar irradiance. This results in a regular and remarkably symmetric pattern with high values occuring near the sub solar zone, ~ 30° north in north summer and ~ 30° south during north winter. The general picture looks similar if separated for land and ocean with the exeption of some interesting features. The north summer net flux maximum occurs nearly at the same time (June) but at higher latitudes (~400) over the land masses than over the oceans ('" 23°). This is not so much the case in the southern hemisphere. The values over oceans have furthermore generally a higher amplitude than over land. Vonder Haar et al. [230, 296J have by subtraction of the net flux at the surface accomplished a representation of the atmospheric energy divergence. This shows marked differences over land and sea. The evaluation implies the existence of a net energy transport by the atmosphere from the continent to the ocean during northern hemisphere summer and the reverse in northern hemisphere winter.

Effects of Changes in the Concentration of Atmospheric Constituents on Energy Fluxes and Surface Temperatures

Climate Research Aspects

The environment on Earth is primarely determined by two factors: the insolation at the top of the atmosphere and the response of the atmosphere-Iand-ocean system to the geographical distribution of this external forcing. The reaction to the solar energy flux depends on the distribution of land and ocean masses, the composition of the surface, and the distribution of chemical constituents (gases, aerosols and clouds) in the atmosphere. The physical state of the environmental system is called the climate, which manifests itself in the energy fluxes and transformations occurring in the system. Any change of properties which determine these fluxes leads potentially to a climate change.

It is known that glaciation periods, which occurred on Earth in the past, correlate very weIl with certain changes in the earth orbit parameters [297-301]. From this knowledge it seems to be justifiable to extrapolate into the future. The conclusion to be derived from the astronomical evidence is, that major climate events, a new glaciation, may not be expected until about 60,000 yr from now. In about 3,000 yr mankind must be prepared for a transitory weaker cooling period, perhaps like the "little ice age" which it experienced from ab out 1,400 a.c. till the nientienth century. At least it seems to be certain that the climate optimum of the present interglacial has passed. The question which is pressing for the near future is, if the decent to a new glaciation period may be accompanied by smaller climate distortions. These could be small in amplitude but nevertheless disastrous for specific regions of the globe, because the dependence on an undisturbed food and energy supply gets more and more critical in a world with growing


populations. Climate excursions in time scales of 10-100 yr can not be excluded, at least not on a regional scale. They have been observed before, and also other major glaciations seem to have been preceded by aperiod of destabilization. In fact regional climate variations have been observed during the last 50 yr in different parts of the world, some of which have been very severe in respect to their impact on food production and water supply [302].

Mankind is presently also facing a new situation. For the first time in history human activity seems to have gained the potential to influence the further development of climate. Processes which could lead to climate modifications are, for instance land surface transformations due to extensive land use which affect surface albedo and hydrology, and the release of substances into the atmosphere which influence its radiative transfer.

Presently it is very difficult to judge possible results of such activities, and to separate the anthropogenic effects from natural variations. Not enough is known about the interactions between the different processes and possible feedback mechanisms, which may either enhance or reduce the initiated trends.

The World Meteorological Organization (WMO) and the International Council of Scientific Unions (ICSU) have therefore agreed upon a joint World Climate Research Programme by which these questions are to be studied for the next decades.

The objectives of this programme are defined as folIows: Our socia! and economic life is vulnerable to periods of climate stress. Human activity may itself

influence local, regional and global climate. These are problems which the international community should address through a World Climate Research Programme (WCRP), which will attempt 10 determine why, how and where climate changes and variations occur, and thereby attempt prediction of their future occurrence.

7he major objectives of a World Research Programme should be 10 determine: - To what extent climate can be predicted. - The extent ofman's influence on climate. To achieve these objectives it is necessary: 1. To improve our knowledge of global and regional climates, their temporal variations, and our under

standing 01 the responsible mechanisms. 2. To assess the evidence for signijicant trends in global and regional climates. 3. To develop and improve physical-mathematical models capable of simulating, and assessing the

predictability of the climate system over a range of space and time scales. 4. To investigate the sensitivity of climate to possible natural and man-made stimuli and 10 estimate the

changes in climate likely 10 result from specijic disturbing influences.

Within this programme so far the following sub-programmes have been recognized [303] as of high relevance to the problem, and implementation plans are worked out for those which have been identified as high priority projects. - Oceanic Processes : large scale circulation, heat transports and heat divergencies,

ocean-air interactions. - Cloudiness and Radiation: investigation of the magnitude of possible feedback

mechanisms on the earth's radiation budget due to changes in cloudiness. - Ocean Cryosphere Processes : feedback of sea ice cover on the radiation budget

by means of albedo changes. - H ydrology and Land Surface Processes : transport of energy due to phase changes

between atmosphere and underlying surface and feedback on the radiation balance due to surface albedo changes, ice and snow cover.

254 H.-J. Bolle

- Radiatively Important Gases: investigation of the effect of the different radiative1y active gaseous constituents on the heat balance of the atmosphere, and c1imatic impact, if the concentrations of these gases change in time.

- Aerosols: assessment of the effect the aerosols have on the radiation budget, and c1imate impact of aerosol concentration changes.

- Solar- Terrestrial Relationships: response of the atmosphere-ocean-Iand system to a long term redistribution (geographically and seasonally) of solar irradiance, and to short term solar variability, e.g. at short wavelengths. There have already been some studies to estimate the significance of certain

atmospheric constituents for the c1imate and the sensitivity of climate for variations in the concentrations of these atmospheric constituents. In conclusion of the discussion on energy transports in the climate system it seems to be appropriate, therefore, to discuss briefly the possible changes in the response of the c1imate system to the incident solar radiant flux due to a variation of constituent concentrations which may partly be induced by man's activities.

Climatic Impacts of Specific Atmospheric Constituents

Clouds

Inbetween all atmospheric constituents the clouds have dominating influence on the radiative transfer of the atmosphere due to their generally high albedo and infrared emissivities. Their effect on the planetary radiation budget can be demonstrated by a simple calculation based upon Eq. (177). The averaged global radiation balance was there defined by

(225)

If N denotes the mean cloud coverage, Qe the cloud albedo, Q' the planetary albedo in cloud-free areas, then Qp can be expressed by

(226)

A similar procedure can be applied to Te. Let Tc be the cloud top temperature and T~ the equivalent te~perature in cloud free regions, then

7;,4 = N7;,4 + (1- N)T:4 . (227)

The temperatures on the r.h.s. of Eq. (272) can be re1ated to the temperature gradient rand the altitudes Ze of the c10ud top respectively z' of the effective emission layer in c10udfree areas:

7;,= T.+rze, T'= T.+rz'. (228)

T. is the surface temperature. With Eqs. (226H228) the balance equation (225) can be written


The Lh.s. has been separated into cloudy and cloudfree areas and the terms can be approximated by the first two terms of the binominal expansions, from which follows

Differentiation results in

or

(132)

Zc and z' can approximately be determined from the balance Eq. (225) if the equilibrium is computed for Qp=Qc and Qp=Q' respectively. For the present state one might use the following numerical values: N =0.5, Q' =0.12, Qc=0.5,

1;=288K, r=-6.5Kkm- 1. Then from Eq. (225) with Qp=Q' respectively Qc:

T' = 270.0 K and '4 = 234.4 K and with Eq. (228): z' = 2.7 km, Zc = 8.3 km. If these numerical va lues are used in Eq. (232) it results with So = 1,370 W m - 2

and u= 5.6696.10- 8 W m- 2 K -4:

~~ ~120K. (233)

Under the condition that the product of atmospheric lapse rate rand the effective emission heights as weil as the difference between the two albedos remain constant or smalI, what means approximately unaffected by cloud cover changes, the temperature change would be 1.2 K per percent cloud cover change.

It can be seen from Eq. (232) that also the height distribution of the clouds (zc) have an effect on the surface temperature, though a change in cloud heights gives a smaller effect than the albedo change [first term Lh.s. in Eq. (232)].

The expression derived in Eq. (232) considers only one cloud layer and disregards the vertical extensions of the clouds, and, even more important, their longwave emission to the earth surface. The model must therefore be refined in order to reflect more realistic the atmospheric conditions. But notwithstanding this draw-backs it represents a zero order approximation to the cloud-radiation feedback problem. The more general questions which have to be solved are: what processes regulate the cloud distribution and what feedback mechanisms are set into action, if the cloud distribution or the cloud cover fraction is changed, and what, finally, is the effect of realistic cloud distributions on the surface temperature.

Attention has been given to the relative importance of the cloud-albedo effect on the one hand and its greenhouse effect due to longwave emission on the other hand. An increase of cloudiness would increase albedo and thus decrease the absorbed solar radiation (albedo effect) but would also decrease the effective emission of the surface if the cloud bases are not at the same time elevated

256 H.-I. Bolle

considerably in height (greenhouse effect). These two effects may therefore cancel each other with respect to the impact on surface temperature, as has been suggested by Cess [304]. In early sensitivity tests prescribed cloudiness has been used instead of internally genera ted clouds and the following results have been obtained [305]:

{

- 0.82 K/% for low clouds

LI T. - 0.39 K/% for middle high clouds 100LlN = +0.37 K/% for infrared black high clouds

+ 0.04 K/% for infrared half black high clouds

(234)

The short wave albedos of the clouds were taken as 0.2, 0.48 and 0.69 for high, middle and low clouds respectively.

In more elaborate studies with GCM'slO Ohring and Clapp [306] concluded that the two competing effects do not cancel each other and that the albedo effect is larger then the greenhouse effect. They found that for the net flux (SW + LW) at the top ofthe atmosphere

acP~ = -57 W -2 aN m ,

resp. - 67 W m - 2 depending on the set of cloud data used, and furt her

acPtw, 00 lacP~ ~o 63 aN aN - .

(235)

(236)

Shukla and Sud [307 a] made experiments with GCM's which generate their own clouds and others in which cloudiness was forced to be constant. It was demonstrated that significant differences in evaporation, precipitation, sensible heat fluxes and radiation fluxes occur, and that the cloud radiation feedback is important in the general circulation of a model atmosphere. Stephens and Webster [307 b] confirmed in general the results of Manabe and Wetherald [305] that low clouds tend to cool and high clouds to warm the surface. A 4% change in low and middle clouds would have the same effect as a 1 % change of the Solar Constant. They could, however, further demonstrate that the magnitude and sign of the effect depends also on the albedo of the underlying surface. There may exist for each cloud type a critical surface albedo below which the cloud will cool the surface and vice versa.

Clouds must therefore be regarded as one of the most critical climate parameters and a reliable cloud statistics is an important prerequisit for climate studies [307 c]. Presently nothing is known about any trends in cloudiness not to talk about cloud top heights variations. Paltridge [308] has estimated that because of the ± 3.5% change in solar irradiance due to the elipticity of the earth orbit the global cloudiness in January should exceed that of July by about 5%. Even this rather strong signal can not be extracted with confidence from existing observations.

It is therefore planned to initiate an International Satellite Cloud Climatology Project (ISCCP) within the WCRP from about 1983 on for at least five years to establish a reliable cloud climatology as a basis for further sensitivity studies.

10 General Circulation Models


Carbon Dioxide

The observed global increase of the e02 concentration in the atmosphere (by about 0.3% per year) is expected to contribute to a rise in surface temperature mainly (by 82%) due to the emission in its infrared bands between 14 and 16 !illl. Due to the increased greenhouse effect the whole troposphere is expected to warm.

This process has been studies extensively by Manabe and collaborators [309, 31OaJ, who used a three dimensional GeM as weIl as a one dimensional radiative convective model. According to their model results, a 100% increase of the e02

concentration in all levels results in a mean temperature increase at the earth surface of about 2 K. The three dimensional model gives slightly higher numbers with a much stronger warming at high latitudes (8 K at 85°, 6 K at 90° and 73°, 4 K at 55°) and about 2 K in the tropics. The area mean temperature increase is in this case 3 K.

In the stratosphere results a temperature decrease from the longwave emission of e02 into space. There is no compensation for this effect due to absorption of solar radiation because e02 has only very weak bands in the shortwave region. The mean additional cooling is about 10 K in 40 km and 5 K in 30 km altitude for doubled amount of eo2. The cooling is stronger in the tropics ( ~ 8 K) than over the poles ( ~ 6 K) at 30 km and about 1 Klarger over the poles than over the equator in the lower stratosphere.

Near the tropopause no temperature changes occur; especially the lower troposphere is heated. Thus the temperature lapse rate in the troposphere would increase by about 0.67 K/km in the polar region and 0.1 K/km in the tropics. The combination of warming in the lower troposphere and the increase of the temperature gradient is expected to have consequences for the evaporation, convection and possible also cloudiness.

Wetherald and Manabe [309J were able to identify a high latitude region at the continents where due to an increase of atmospheric e02 the precipitation and runoff rate increases markedly, a zone of decreasing soil moisture around 42° latitude, and a zone of enhanced wetness in subtropical regions. Further results are a general warming, an increase in moisture content of the air, a reduction of the meridional temperature gradient in the lower troposphere, a poleward retreat ofthe snow boundary and an increase in the polward transport of heat. With this model also the two dimensional (height-Iatitude) change in cloudiness could be studied more detailed as previously possible. With increasing e02 cloudiness decreases in the upper and middle atmosphere at most latitudes. The low cloudiness, however, increased remarkably at high latitudes and in the sub-tropics in the lower troposphere. An increase in cloudiness was also recognized in the lower stratosphere of the model.

Sensitivity studies ofthe e02 impact on climate have also to include the oceans which store and release eo2, depending on their temperatures. Only few attempts have so far been made to include the oceans in such models [31Oa, b].

Numerical model studies are generally carried out with a two- or fourfold e02 concentration. Between 1900 and 1980 the e02 increase is in the order of 10%. The signals to be expected right now are therefore much smaller than in the model studies. No temperature trend has so far been observed which

258 H.-J. Bolle

can easily be correlated with the CO z increase [311a]. The signal would still be very small compared with the "noise" of the temperature fluctuations. Hansen et al. argue in a very recent publication [311 b] that the reason for the COz temperature effect not being evident in the records is its present compensation by volcanic aerosols (Table 29) and variations in the solar radiant energy output. They claim that if these superimposed effects are accounted for a warming of the troposphere is evident in the right order of magnitude (about 0.3 K during the last century).

Table 29. Surface temperature effect ofvarious global radiative perturbations, based on onedimentional radiative-convective model computations by Hansen et al. [311 b]

Atmospheric Quantity Change ,1 T, K

CO2 300 ppm->600 ppm +2.8 Solar irradiance +1% +1.6 High clouds +2% ofglobe +0.9 N20 0.28 ppm->0.56 ppm +0.6 Tropospheric aerosol; soot ,16= +0.02 +0.5 CCl2 F2 and CChF 0-+2 ppb each +0.5 CH4 1.6 ppm-> 3.2 ppm +0.2 Middle clouds +2% ofglobe -0.4 Ozone -25% -0.5 Tropospheric aerosols, H2S04 ,16=+0.1 -1.2 Land albedo +0.05 -1.3 Low clouds +2% ofglobe -1.4 Stratospheric aerosols, H2S04 ,16= +0.2 -1.9

The temperature signal of the COz increase should be less disturbed and maximum in the tropical middle stratosphere where it may eventually be picked up by satellites or high altitude ballon so undings. A cooling in the stratosphere would have a side effect on the chemistry of the stratosphere because the reaction rates are temperature dependent. This could affect the distribution of stratospheric trace constituents.

Ozone

According to Ramanathan and Coakley [312] the surface temperature on earth is equally sensitive to a change ofthe ozone concentration as to a vertical re-distribution ofthe ozone mass. The application oftheir radiative convective model indicates that reduction of the ozone concentration in all heights will produce a change of the earth surface temperature of -0.014 K per unit percentage change in ozone concentration if the cloud top temperatures remain fixed, and - 0.009 K if the cloud top altitude is fixed - all other parameters remaining unchanged. Wang et al. [313] gets under the same conditions 0.019 K and 0.014 K respectively. An equally large effect as a 50% decrease in the ozone concentration can be achieved by a 4 km upwards shift of the maximum of the ozone distribution: a re duc ti on of the temperature at the surface by 0.44 K. By the upward shift of the


radiating ozone mass the emission lines are narrowing proportional to the pressure which reduce the emissivity. With a general circulation model Manabe and Weatherald [305] computed earlier a temperature change of 1.4 K if the ozone amount increases by 0.1 cm STP ozone or of ab out 30%. This would result in 0.047 K per % change of the ozone concentration which seems to be rather high in the light of the more recent results of the other authors.

Other Minor Trace Gases

Much attention has been paid during the last decade to those trace gases which are potentially responsible for changes in the ozone mass budget by means of photochemical reactions. Among these are species which themselves have strong absorption bands and could therefore in the course of further accumulation in the atmosphere become important for changes in the radiation budget. Sensitivity studies of the effect of these molecules on the change of the earth's surface temperature have so far been made with rather simple models in which a number of feedback mechanisms are not included. The computations indicate, however, which of the constituents have to be included in radiation budget studies if their concentrations in the atmosphere exceed certain limits.

The largest effects can be expected from N2 0, NH 3 , CH4 , and HN0 3. These are naturally controlled species wh ich have so far been considered as fairly constant in the atmosphere. N2 0 is produced by fertilizers, however, and its extended use could increase its concentration by a factor of2 even within 50 yr. This would then in turn also affect the concentrations of NH3 and HN03 . Another category are methane derivates such as the chlorofluorcarbons which contribute negligibly to the radiation budget at the time being, but their concentrations increase much more rapidly than those of the before mentioned species and of CO2 . It can therefore not be excluded that their effect will exceed that of the other species in the future.

Of all minor constitutents a change of the average tropospheric water vapor concentration would have tremendous impact on climate. The GeM computations ofManabe and Weatherald resulted in a 0.13 K increase in surface temperature if the humidity is increased by 1 %.

Wang et al. [313] reported temperature changes of 1 K for the fixed cloud top temperature model and 0.65 K for the fixed cloud height model, if only the amount of stratospheric water vapor is doubled above 11 km while the tropospheric relative humidity remains fixed. Similar results were also obtained by Manabe and Weatherald. According to their model a 5-fold increase of stratospheric water vapor mixing ratio from 3 . 10- 6 to 15·10- 6 would result in a 2 K increase of the surface temperature.

The effect of an increase of the trace species on the earth's surface temperature as climate sensitivity parameter are listed in Table 30.

The direct effect which an increase of N02 would have in the climate system is not yet adequately evaluated. N02 reaches concentrations up to 10 ppbv .with a large diurnal variation in the stratosphere. Nitrogen dioxide absorbs in the shortwave region between 300 and 700 nm with peak values between 410 and 460 nm (Leighton [314], Dixon [315], and Hall and Blacet [316]). In the infra red

260 H.-J. Bolle

Table 30. Increase LlTo of the Earth surface temperature per percent increase of minor trace constituent concentrations after [312a-c, 313]

Species

H20 H 2O' CO2 0 3

N20 CH4

NH3

HN03

S02 C2H 4

CCIz CHCl3

CH2Cl2 CH3Cl CCl4

CF2Clz CFCl3

• Stratosphere only

Assumed present concentration C (ppmv)b

3 .10- 6

330 3.43mm STP 0.28 1.6 6 .10- 3

4.87.10- 3 mm STP 2 .10- 3

2 .10- 4

1 .10- 4

1 .10- 4

1 .10- 4

5 .10- 4 } 1 .10- 4

1 .10- 4 } 1 .10- 4

b 0 3 and HN03 : colurnn amount e Absorption data probably too low

LlTojO.01 C fixed c10ud top

tempo height

in 10- 3 K

10 32

-19 6.8 2.8 e

1.2 0.8 0.3 0.1 0.15 0.1 0.01

0.2

0.3

130 6

21 -14

4.4 2.0 e

0.9 0.6 0.2 0.1

0.1

0.2

Valid for concentrations around ppmv

average humiditv 6-15.10- 6

412 2.57 mm STP 0.56 3.2

12 .10- 3

9.74.10- 3 mm STP 4 .10- 3

4 .10- 4

2 .10- 4

1 . 10- 3

1 . 10- 3

{ I .10- 3

2 .10- 4

{ 2 .10- 3

2 .10- 3

there are two strong vibration rotation bands at 648 cm - 1 and 1,621 cm- 1

(Herzberg [37]). Both are imbedded in the strong atmospheric bands of carbon dioxide and water vapor respectively. Zdunkowsky et al. [317] have included the shortwave band of NOz in a planetary boundary layer model. The effect on atmospheric and surface temperatures is small unless the concentration reaches 1 ppm.

Table 30 indicates that there exist a number of trace gases which at least in their totality are equally important as COz. This is especially true since the accumulation in the atmosphere of some of them occurs at a higher rate than for COz. It has to be considered that the cumulative impact of these trace gases on the greenhouse effect may grow to 2 K between now and the first decades of the next century.

Complete1y unsolved at this time is the question whether the water vapor amount in the troposphere remains constant or not. It would be necessary to monitor the global water vapor budget to aprecision of at least a few percent per year in order ot detect trends within a decade.

A more accurate assessment ofthe magnitude ofthe effect the trace gases have on the energy fluxes in the atmosphere requires a more complete world wide monitoring system oftrace gas concentrations, and sensitivity studies with refined numerical models (compare final section).


Aeroso[s

General

261

Aerosols scatter and absorb radiation at short wavelengths which reduces the amount of solar energy available at the surface for heating and evaporation. A purely scattering aerosol enhances albedo and will thus decrease the amount of energy available in the atmosphere-earth system. As the aerosol starts to absorb it will even more reduce the irradiance at the surface but the absorbed energy will directly be transferred to the atmosphere (instead via the surface and upward heat transport). The amount deposited in the atmosphere will be larger over high albedo terrain than over low albedo surfaces because in that case more radiant energy gets the chance to travel through the atmosphere and getting absorbed twice. Eventually the heat gained by the atmosphere may even exceed the losses at the surface and the net effect may become a warming of the system.

At wavelengths longer than 3).ill1, in the infrared, absorbing aerosols will emit radiation to space and extract this energy out of the atmosphere. It will also reduce the effective radiation of the surface and contribute to the greenhouse effect.

The resulting net effect depends on the relative magnitude of the optical properties of the aerosol in these two wavelength regions, and on the ratio of absorption to backscatter. For each value of this ratio exists an equilibrium value ofthe surface albedo for which neither warming nor cooling occurs. For maritime aerosols these equilibrium albedo is high (0.6 ... 0.9), for continental aerosol it ranges between 0.1 and 0.6. Over high albedo surfaces therefore nearly all aerosols cause warming, and over the sea surface only the most absorbing aerosols (carbonaceous particles) have a warming effect. An assessment of the aerosol effects on global climate is therefore equally important [318aJ as those of other atmospheric constituents.

Sensitivity Studies for Tropospheric Aerosol

SeIlers [318 b J estimated that presently to the optical depth of the atmosphere (0.3) molecular scattering is contributing 0.145 and particle scattering 0.155. Of the particle scattering optical depth the amount of 0.023 is due to man made aerosols. If this amount is doubled the computed surface temperature change is -0.5 K and if the man made aerosol amount is tribled the temperature drop would even be - 3.2 K. This is a very rough estimate in which only the optical depth effect, not the absorption is taken into account. The real effect will probably be smaller.

Since the first studies of SeIlers more detailed aerosol models have been explored. One of the findings [319, 320J is that the imaginary part of the refractive index and the optical depth b are more critical parameters than the phase function and the mean radius r. Hansen et al. [321J found for Wo a critical value 0.85 above which cooling and below which warming occurs. In regions with strong industrial concentrations therefore warming can be expected while in other regions the aerosol may cool if the imaginary part of its refractive index in the shortwave range is low. The temperature effect may range between + 3 K (wo =0.6) and -1 K (wo = 0.95) [321]. A complete omission of aerosol in a zonally

262 H.-J. Bolle

averaged model would introduce an error of 1 K in the computed surface temperatures and an increase of the optical aerosol depth would in such a model decrease the surface temperature by [322]

LlT/Llc5=-11.6K.

Additional consideration of doud albedo feedback may, however, reduce the temperature effect [323]. Other results of Ohring's work are that in an atmosphere with constant relative humidity the effect on surface temperature is stronger than in an atmosphere with fixed absolute humidity.

Charlock and SeIlers [324] computed the aerosol effect with a time dependent monthly as weIl as an annual model. By adding aerosol according to a model of Toon and Pollack [325] with an optical depth of 0.125 the surface temperature is reduced by 1.6 K or dT/dc5 = -12.8 K, a value which is dose to that one found by Ohring. For a strongly absorbing aerosol, wo=0.75, their annual model gives dT/dc5= +4.8 K at 40-50° N latitude with a surface albedo of 0.139 while no temperature change occurs for Wo ;::;:0.81.

The dependence of the aerosol effect on the surface albedo has been studied by Reck [326, 327]. If the atmosphere would be optically dense with an optical depth of c5 vis ~ 80 in the visible, cooling would occur ifthe surface albedo Qs is < 0.35, and heating if it is larger. In an optically thin atmosphere (c5 vis ~4) the neutral albedo value is Qs = 0.6. F or intermediate values of the optical depth, 4< c5 vis < 80, there is a transition region where the line for zero effect changes from Qs=0.6 tOQs=0.35. Thus under normal conditions with small optical depthand low surface albedo, there will always be cooling.

From a comparison of computations with [328, 329] and without [322, 330] infrared aerosol absorptivity it results that the main effect of tropospheric aerosol is due to backscattering of the short wave radiation flux, a result which was first recognized by Rasool and Schneider [329]: The bulk of the aerosol is concentrated in the lower atmosphere where the temperatures are not much different from surface temperatures, especially if one considers daily averages with the nocturnal inversions.

Over deserts, however, the situation is different. Here it can occur that silicate material is tran spor ted high up into the troposphere with an imaginary part of the refractive index three to five times smaller than und er undisturbed conditions [331]. In these cases the infrared radiation properties gain in importance: the reduced local heating due to absorption of solar radiation is partly compensated by the infra red emission. At the surface heating seems to prevail: the deficit in solar energy reaching the ground is partly compensated by the longwave opacity and its greenhouse effect.

In cases of dust outbreaks from the Sahara over the Atlantic as studied by Carlson and Benjamin [332] for the GA TE area aerosol was injected up to tropospheric levels around 500 mb. Infrared cooling rates up to - 3 K/day resulted in the center of the dust doud far a total optical depth at 550 nm around 3.0. But also in these cases the net effect averaged from 1,000 mb to 500 mb and 24 h was a warming by about 1 K due to the combined effect: net absorption of solar radiation in the dust layer and an additional greenhouse effect in a


layer between 1,000 mb and 900 mb due to the infra red aerosol properties. If the same aerosol is over the desert, the greenhouse effect is converted into cooling near the surface but still net heating is maintained if the whole layer is considered.

Sensitivity Studies for Stratospheric Aerosol

Stratospheric aerosol is genera ted by both, gas-to-particle conversions and volcanic events. It therefore consists of two components, one permanent one and another compenent which is injected locally, spreads globally by turbulent mixing and has a residence time of a few years. Thus the scales for the aerosol in the stratosphere are quite different from those observed in the troposphere. On the other hand the optical properties of the stratospheric aerosol seems to be more uniform than of that in the troposphere. However, on a closer examination this may be a wrong impression with respect to the volcanic component since there are indications that each volcano may genera te its specific aerosol. The climate effect can again be computed from a data set of Wo, < cos 9), r, 6 (vis) and the albedo of the underlying levels which is very close to the planetary albedo (lp.

Typical values for the optical aerosol parameters are given for instance by Shettle and Fenn [333].

The heating rates in the stratosphere turn out to be very sensitive for a change in Wo [337,335]. Va lues of 1-1.5 K/day have been computed at 17 km for Wo between 0.6 and 0.8. In the troposphere a stratospheric aerosol layer can generate a warming (in the order of 0.1 K/day) if wo< 0.98 or a small cooling effect if Wo is approaching 1.0. The magnitude of the tropospheric cooling depends on the infrared value of Wo, being largest if there is no infrared opacity.

Due to the stratospheric background aerosol the spherical albedo of the earth can be changed only in the order of less than 10- 3. For single scattering albedos larger than Wo = 0.94 the aerosol will increase the earth albedo and below this critical value it will decrease it. For climate research it is therefore extremely important to determine very accurately the albedo for single scattering in the region just above 0.9.

Feedback Processes with Aerosol Involvement

The aerosol-temperature effect can not be seen isolated from the behavior of the whole climate system. Little is known so far about the magnitude of possible feedback mechanisms. But it has been mentioned earlier that another major aspect of aerosol physics is the ability of particles to function as condensation nuclei. In addition aerosols may be imbedded in clouds inbetween the cloud droplets, thus affecting again the optical properties of the clouds [338]. Another cloud related effect is the dissolution of clouds in areas of desert aerosol outbreaks over the oceans, a phenomenon which can often be observed in satellite images. By this effect the planetary albedo may rather be decreased in the areas affected by the desert aerosol then increased as assumed in some simulation studies.

The combined effects of climate sensitive parameters on the temperature field may result in a change of horizontal (latitudinal) temperature gradients which would change the general circulation and heat transports. Also the stability of the

264 H.-I. Bolle

atmosphere depends inter alia on the radiation effects of the aerosol. Such processes may lead to changed precipitation patterns and wind speeds. Joseph [339a] has demonstrated such effects with a general circulation model. If arid zones are affected in such a way that its climate becomes dryer and that from time to time more intense storms develop, then also the amount of aerosol may increase over these regions and be transported over large distances thus enhancing the aerosol effect (compare also [339b]). A look at Mars demonstrates the extrem possibilities which exist on a dry planet. There are also indications [339c] that during the last ice age the terrestrial atmosphere was much heavier loaded with aerosols than it is nowadays.

Anthropogenie pollution is another mechanism which, because of its permanency, may already have affected climate in certain regions. On aglobel scale this effect can still only be small because of the short residence time of these aerosols, which are not transported high up into the atmosphere. There is, however, a man induced feedback mechanism in arid zones. Extensive overgrazing and agriculture enforces erosion in these regions which again increases the probability that dust is raised from the ground and mixed up into the middle troposphere thus adding aerosol. The surface albedo change connected with such

Table 31. Vo1canic eruptions 1958-1978 and maximum injection heights. (After Herbert [340])

Year Volcano Injection height

1962 Mt. Trident 12km DeFuego 12

1963 Agung 20 km Mt. Trident (2 x) 12

1964 Irazu 12 Sertsey <10 Lopevi 10

1965 Taal 15 1966 Redoubt 13

Kelut <10 Soputan <10 Awu <10 Taal <10

1968 Femandina >20km 1970 Hekla 15

Beerenberg (2 x ) 12 Deception Is. <10

1971 Mt. Hudson 10 DeFuego 15 Artcic Circle 15

1972 Ambrym 10 1974 DeFuego 20 km 1976 St. Augustine 10 1977 Nyiragongo 10

Mt. Usu 12 1978 Westdahl 10


an erosion will again act as a feedback mechanism by reducing the tendency for cooling of the whole system with increasing surface albedo [327].

R6le of Vo1canism

Volcanic eruptions eject large amounts of solid particles, ashes, as weIl as gases into the atmosphere. Both components are able to affect the radiative transfer in the atmosphere. The particles scatter, absorb and emit radiation and the gases interact with the radiation field if they have absorption bands. As a secondary effect some gases like S02 form aerosol particles consisting of H2S04 in the stra tosphere.

If during an eruption volcanic material is direct1y injected into the stratosphere, the smaller particles « 1 j..Ull) can remain there for a long time. Particles of a density of 1 g cm - 3 and a size of 1 j..Ull radius have a sedimentation velo city of about 2.10- 4 m S-1 at 10 km altitude and particles of 0.1 j..Ull radius of 6.10- 6 m S-1. Particles smaller than 1 j..Ull can therefore stay at altitudes > 10 km for in the order of a year and particles smaller than 0.1 j..Ull for decades.

Most of the explosive volcanous eruptions remain below 10 km (Table 31). Only after the eruptions of Agung (1963), Awu (1966), and DeFuego (1974) major

Table 32. Density ofthe different aerosol components at the Wank-Mountain under extreme clean air conditions above an inversion with an austausch coefficient smaller than 5 g cm - 1 S -1 and composition of Sahara Dust. (After Kunkel et al. [340] and Reiter [341].) The number o[ condensation nuclei was 313 cm- 3 at the Wank station (1780m) Wank station:

Aerosol components

Na+ K+ Ca2 +

CaO insoluble Fe203 Si02

Sahara dust:

Component

Si02 AI20 3

CaC03 Fe203 Mn3 04 MgO

Conc. in n g m- 3

56 33 63 86 66

353

Relative concentration

37 -75% 0 -20% 1 -10% 6 -22% 2 -4% 0.4- 3%

Traces of K, Cu, H 2S04 , HCI Size distribution: 68 % r ~ 0.5 fIIl1

24% 0.5<r< I fIIl1 4% r;;; 1 fIIl1.

Aerosol components

Ah03 Pb2 +

CISoiN03

NH4

The Fe- and Mn-oxides content in the dust depends on local sources

Conc. in n g m- 3

163 10 76

546 412 351

266 H.-I. Bolle

sharp decreases of the atmospheric transmission have been observed, a small decrease eventually after Fernandina (1968). From 1962 to 1965 the atmospheric transmissivity dropped by 0.018 (from 0.933) at Mauna Loa.

Some data on composition of non-volcanic aerosols are compiled in Table 32 which is adopted from Kunkel et al. [341 J and Reiter [340J, Table 33 a from [342J, and Table 33 b after Zoller et al. [349]. In contrast to these background values the analyses offumes from craters are given in Table 34. These emissions vary strongly with time; thus the values given are representative only for the time the sampIes were taken.

Table 33a. Mean atmospheric chemical concentrations of atmospheric aerosol at South Pole Station and Mauna Loa after Herbert [342]. For As and Sb the results from two methods are presented (Nuclepore and Whatman 541 filters have been used)

Element Unit South Pole (all data) Mauna Loa (clean)

Summer" Winter b Downslope

S ng m- 3 76 ±24 29 ±1O 74 ±12 Na nm- 3 5.1 ± 1.7 40 ±31 3.3± 2.5 Al ng m- 3 0.83± 0.41 0.30± 0.04 5.0± 3.2 Mn pgm- 3 14 ± 6 6.7 ± 4.5 85 ±65 V pg m- 3 1.6 ± 0.6 0.9 11 ± 9 As pgm- 3 24 ± 7 17 ± 3 As pgm- 3 8.4 ± 1.1 17 ± 9 Se pg m- 3 6.3 ± 0.6 6.9 ± 2.7 35 ±24 Sb pg m- 3 2.0 ± 1.6 6.6± 2.9 Sb pgm- 3 0.45± 0.16 2.1 ± 1.5

" Average of four summers b Average oftwo winters

Table 33 b. Mean atmospheric concentrations of trace elements at South Pole Station. (After Zoller et al. [349J)

Element Mean concentration" Element Mean concentration"

Na 7.2 ± 3.8 ngm- 3 Se 5.6 ± 1.2 pgm- 3

Mg 1.0 ± 0.7 ngm- 3 Cr 5.3 ± 3.0 pgm- 3

Fe 0.84± 0.21ngm- 3 Ce 2.3 ± 1.6 pgm- 3

Pb 0.63± 0.30 ng m- 3 Sb 1.7 ± 0.6 pgm- 3

Br 0.63± 0.30 ng m- 3 V 1.5± 0.6 pg m- 3

Al 0.57± 0.17 ng m- 3 Co 0.84± 0.27 pgm- 3

Ca 0.5 ± 0.4 ngm- 3 La 0.51± 0.37pgm- 3

K 0.3 ± 0.1 ngm- 3 Sc 125 ±48 fgm- 3

Cu 36 ±19 pgm- 3 Th 59 ±21 fgm- 3

Zn 30 ±11 pgm- 3 Sm 55 ±23 fgm- 3

Mn 10.35± 5.5 pgm- 3 Eu 17 ± 4 fgm- 3

" The uncertainty reported in the standard deviation of the mean concentration in 10 sampling periods

Upslope

230 ± 90 88 ± 56 14 ± 8

190 ±11O 27 ± 13 23 ± 20

25 ± 8 5.8± 3.3


Table 34. Analysis of sampIe of particles eolleeted from the fume of the lava lake of Mauna Ulu erater of Kilauea for 15 min at 15 I per min on polystyrene fiber filters on Oetober 19, 1972 (top left) and of particles eolleeted from the fume from HaIemaumau erater of Kilauea on April 5, 1972 for 5.0 h at 15 I per min on polystyrene fiber filters (top middle and right). (After Cadle et al. [343 al) Composition offumarolie gases from Showa-shinzan, Japan (boltom). (After White and Waring [343 b])

Element Weight collected on Element Nanograms eollected Element Nanograms colleeted orion Filter (Nanograms) orion on Filter orion on filter

SO~- 208,000 SO~- 4.1 . 106 Cr 200 NHt 1,200 NHt 4,500 Sb 200 CI 2,700 CI 4,800 Fe 4,000 Na 23,000 K 0 Co 20 K 0 Na 500 Sc 4 Mg 1,100 Mg 1,500 Ni 0.0 Pb 180 Ca 22,000 Cu 900 Ca 3,600 Si 16,400 Cd 37 Cu 550 Mn 250 Pb 125 Ni 0 Br 140 As 24 Cd 180 Hg+Se a 80

a About 50% Hg

Elementor Parts per million by Element or Parts per million by molceule weight at moleeule weight at

760°C 220 D C 760°C 220°C

CO2 29,200 13,000 B 39 6 CO 50 P04 2.8 1 S02 1,490 716 N02 0.01 0.01 S03 21 3 O2 51 23 HzS 8 1,080 Hz 685 20 S 4 NH 3 1 17 CI 728 433 Nz 567 1,250 F 238 35 40Ar 1 1 Br 1 1 CH4 2

Oz and 4°Ar are very probably not released by the voicano, HzO, an additional major eonstituent of voleanic outgassing, is not listed

The eruptions of Mount St. Helens in May 1980 were very carefully studied by aircraft measurements. After several smaller eruptions the main explosion occured on May 18, 1980 at 15:32h, major primary components of the emission were SOz and reduced sulfur (probably HzS) [344] with an emission rate of 1,000-1,500 t SOz per day during the main eruption phase. Also carbonyl sulfide, OCS, and halogen compounds were analysed. One day after the main eruption at an altitude of 14 km the following maximum gas concentrations have been observed [345, 346]:

SOz: 111±24 ppb (0.04-0.05 ppb) CH3CI: 3±0.4 ppb OCS: 0.87 ±0.11 ppb (0.1 ppb)

Cl- (acidic): 2.4 ppb SO~-: 174± 17 ppb N03": 6.2±0.6 ppb.

268 H.-J. Bolle

The pre-eruption mean valuesbetween 15 and 21 km are given in parenthesis. On May 22 maximum concentrations have been found at 20.5 km altitude: CH3CI: 5 ± 1.2 ppb OCS: 1.2 ± 0.35 ppb S02: 0.08±0.04-4.2±0.55 ppb

Especially the concentration of S02 was found to be highly variable. The amount of water vapor also increased considerably fromin the mean of

1-2 ppm to 15-40 ppm at 18.6-20.1 km altitude and 2-25 ppm around 20.1-20.7 km altitude [347].

Table 35. Concentrations and relative enrichments in plume sampIes of Mount St. Helens from 19 May 1980. (After Vossler et al. [348])

Element

B Na Mg Al Si S Cl K Ca Sc Ti V Mn Fe Co Zn Ga As Se Rb Ag Cd Sb Ba La Ce Sm Eu Gd Tb Yb Lu Hf Ta Th

Concentration (/lg/m3)

17.7 km

0.051 96

260 1,100

5.8 2.2

46 130

0.035 16 0.18 1.3

110 0.036 0.39 0.74 0.0043 0.00077 0.14 0.0019 0.099 0.00092 1.3 0.046 0.11 0.012 0.0039 0.014 0.0019 0.0038 00087 0.014 0.0014 0.0031

13.1 km

0.068 79 29

220 840

5.6 3.5

41 56 0.019 8.2 0.082 0.90

72 0.017 0.22

0.026 0.00072

- 0.15

0.00015 0.00105 1.10 0.040 0.079 0.0080 0.0024 0.0074 0.0013 0.0033 0.00049 0.012 0.00093 0.0039

([Element]/[Al])plume

([Element]/[ Al]).sh

17.7 km

1.10 0.97

1.00 1.26 2.35 0.72 1.22 1.52 1.60 1.56 1.38 0.94 1.46 0.89 2.8 1.43 0.79 1.00 1.19

360 2.63 1.18

0.99 1.32 0.91 1.73 1.32 0.84

1.16 1.18 0.51

13.1 km

1.71 0.93 1.37 1.00 1.13 2.70 1.37 1.27 . 0.76 0.97 0.95 0.72 0.74 0.96 0.50 1.82

5.6 1.08 1.52

6.4 3.52 1.18

0.87 1.04 1.17 1.08 1.57 0.87

1.16 0.92 0.75


The element concentrations are given in Table 35, which may be compared with pure air conditions given in the previons tables. The mineral components were [350]: Glass (",60%), Hornblende ('" 10%), and Pyroxene ('" 10%).

The ash injected into the stratosphere had a rather small imaginary part of the refractive index, ",0.01 at 300 nm and between 0.0015 and 0.004 at 700 nm [351]. In comparison the DeFuego ash was much darker, 0.1 at 300 nm and 0.015 at 700 nrn. There seems therefore to be a broad spectrum of variability between the debris of different volcanic eruptions. The size distribution of the particles measured between 12 and 18 km altitude [350, 352] peak around 0.5 ).IID

radius with maximum dN/dr= 5 cm - 3 after 4 days in 18 km altitude and dN/dr",0.5 cm - 3 after 16 days in 12 km altitude.

The distribution of the Mount St. Helens plume in the stratosphere was monitored by satellites (SAM II and SAGE). The analysis of this unique set of data will not only provide global concentration measures but also deeper insight into transportation and mixing mechanisms in the stratosphere [353-355b].

The impact of these eruptions on climate and climate variability can at the time being only roughly be estimated or modelled. The effects of the globally distributed and diluted debris of one volcano eruption is in most cases small. However, as Pollack et al. [335] pointed out, frequent multiple explosions of volcanos could account for a depression in surface temperatures. Periods of frequent explosions occurred e.g. at the end of the 19th century (1882-1890: Krakatau, 80 S; Falcon Is., 200 S; Farawera, 30SS; Niafu, 16°S; Bondai san, 38°N; Ritter Is., 5SS; Bogoslof, 54°N), and a weaker sequence between 1902 and 1914 (Soufriere, 13SN; Santa Maria, 14°N; Shtynbelya Sopka, 62°N; Katmai, 58°N). Also the present activity can be regarded as clustering of events (Table 31).

After strong volcanic eruptions the stratosphere may maintain an optical depth of 0.1 and even 0.3 for aperiod of one or two years [334]. Charlock and SeIlers [324] and Hansen et al. [336,311 b] computed the effect of a volcanic dust layer with time dependent models for the time following the Agung (8 0 S, 115°E) eruption in 1962. For the surface temperatures in the tropics adepression in the order ofO.4-0.55 K was computed for aperiod of about 2-3 yr which is in rather good agreement with an observed depression of 0.5 K with respect to March 1963 lasting from mid 1964 to 1966. In the stratosphere a warming up to 6 K has been observed from autumn 1963 to spring 1964 over Australia which agrees with the models which give the maximum of the warming in la te autumn 1963 but then a steeper decrease to "normal" conditions then has been measured.

Harshvardhan [337] investigated the effect of a stratospheric aerosollayer and got strongest effects in polar regions with very small effects at latitudes < 50°. His conclusion is that perturbations in the order of b ~0.1 as occur after volcanic eruptions may induce climatic significant changes due to the change of the equator to pole radiant gradient.

During periods of high plutonic activity the aerosol concentration in the stratosphere may be maintained in a highly disturbed state over quite long periodes in the order of decades. Simulation results indicate that temperature excursion can be expected in the troposphere which are in the same order of magnitude as generated by the increase of CO2 during the last hundred years: a few tenth of a degree [335].

1.0. -- ice free ----- ice covered

0.9

0.8 r .c. C 0 0..7 E

N I

E 0.6 -, (!)

0..5 ~ :::J 111

0..4 0 Co X C10

c 0.3

" "0

" 0..2 c::

0..1

0.

-0..1

I .c. C 0 E

1: 0..1 -, ~

X :::J

>- -0..1 ~ C10 C w

-0..3

-0..4

-0.5

Month

Fig. 38. Heat budget of ice or snow coverd and ice-free polar areas. Left: Energy budget of areas in the Barent sea (750 N). (After Badgley and Duronin [359] and Marskunova [245]). Upper part: radiation budget components for ice-free (fulI line) and ice-covered areas. Though the actualIy incident solar radiation is much smalIer for the ice-free area (due to c1oudiness), the absorbed solar radiation and the radiation balance are much larger. The radiation budget of the surface is negative ( - 10.9 MJ m - 2 a - 1) for

1.0

0.8

0.5 ., :S c 0.4 0

E N 0.2 ,

E ,

0 <.)

x -0.2 :::J

LL

-04

-0.6

-0.8

100

o Computed . Observed

Q~

/ "'-/ Direct solar flux

/ ~ 0\ ,",10" ""0"'''')

/ \~ r/) ~ ___

Ö~cr--ö~rface longwave emissio~6

~~ ,.. m

/ "'" m* "'SH

"= ,/ ' ...... ~R /.-.--............. c .--e __ .____ ;' ........... •

NE 50 ."", / .~"~. 'E • ,/ .... , \

..... ----. . \ , O~--~~--~~~------~~~~--------~~~------~~~._----~ L ." / ............. ~ ...... __ .~._._. ~ \ /e--.. "

X :::J -50

LL

L

0. (lJ

o

-100

SFC

2

6

..... /' ~ _. , 'l''' \ /' ... ---..... . .....,.... .- , .".,..; '. .

./ ,,/ -", ~'~~'-_. "'./ .., / \ .-, // ....... _-...... - ......... . ............................................ '._-......... / .. .. ...................... --_. .. ....... , .......... ..

tPLH

A S o N D F M A M Month

the ice covered area and positive (825 MJ m - 2 a - 1) for the free ocean. Lower part: Estimated energy budget components for the ice-free artic ocean. The flux between the surface and deeper levels is nearly zero in the annual mean (21 MJ m - 2 a -1). Right: Energy budget parameters for the Antarktic station Mirny. (After Schlatter [246]). Top: Radiation budget for snow covered area. The magnitudes of the components are comparable with those of the arctic ice-covered case. Middle: Energy budget components for the snow covered surface. Bottom: Temperature distribution in the top 6 m of the snow layer. Note the quite different behaviour of the latent and sensible heat fluxes as compared to the icefree ocean (left). The small positive energy balance in December and January causes a ~arming to zero degree which penetrates to 6 m depth with a phase lag of about 2 months

272 H.-J. Bolle

Surface Albedo Feedback

One of the most dramatic changes which occur at the earth surface is the annual variability of the surface albedo at latitudes > 50°. The snow and ice cover of vast land areas and the ice of the oceans modulate strongly the energy fluxes between the ocean respectively the soil and the atmosphere. This affects not only the albedo (and by this the radiation budget) but also thermal conductivity and evaporation. As far as such changes are periodic without an interannual residuum they belong to the normal variability of the balanced state of climate. Ifthere are trends, such as the ice not retreating to its starting position every year but extending further from year to year, then a positive feedback occurs which could lead to rapid climate changes. Positive feedback in this context means that changes introduced by the growing ice mass to the energy transfer system have tre effect of accelerating the generation of more ice rather than to stop the development. In a simple energy balance climate model Budyko [356J has demonstrated that very rapidly a glaciation of the whole earth would occur, if due to reduced insolation the ice cover would proceed to 50° latitude.

This problem has been addressed especiaIly by Lian and Cess [357J, who computed the sensitivity parameter

-

ß=S dTs o dSo

(237)

(So being the Solar Constant) and showed that ice-albedo feedback would enhance this quantity from 147°C to 184°C or by 25%. Energy balance models show that for the present Solar Constant both states can exist, the present climate and a completely ice covered earth, as weIl as an intermediate stage [358]. The present climate is stable against internal changes but not against external changes of the Solar Constant: Climate turns over to the completely glaciated stage if the Solar Constant is reduced by a certain not yet very precisely defined percentage.

The changes which occur if glaciation takes pI ace have been studied in some detail though they could not yet be implemented to a fuIl extent in simulation models. The annual heat budget of ice and snow surfaces versus an open lead have been estimated by Badgley and Duronin, Table 36. The monthly variation of its components is given in Fig. 38.

Table 36. Annual heat budget of iee and snow surfaees versus an open lead at 80° N. (After Fleteher [359].) In MJ m - 2 a - 1

Net solar radiant exposure (absorbed solar energy) Net longwave radiant exposure (negative ifupward) Net radiant exposure (negative ifupward) Turbulent sensible heat flux (negative if upward) Latent heat flux (negative if upward) Conduetive heat flux to the surfaee (positive if upward)

lee and Open lead surrounded snow surfaee by sea iee

+730 -758

- 28} + 40 -15 -27 + 15

+ 2,746 - 3,022

- 276} -10,649 -12,040 - 1,118 +12,040


Table 37. Heat flux through a 2 m thick arctic pack ice layer and from open water surface in W m - 2.

(After Untersteiner [360] and Badgley [361], compare [359])

Surface/layer Oct

Pack ice 0.84 Open ocean 312

Ratio 372

Nov

4.6 565

123

Dec

8.4 800

96

Jan

10.0 800

80

Feb

10.9 887

82

Mar

10.5 862

82

Source

360 361

The ratio of heat flux from an open water surface and through pack iee varies from nearly 400 in Oetober to about 80 in early spring (Table 37).

A more complete assessment of the differences in energy fluxes through the surfaee in different arctie and antaretie loeations (and also eomparisons with model ealculations) have recently been made by Herman and Johnson [362] and also by Weller [363, 373]. Some results ofthese eompilations are reproduced in Tables 38 and 39 respeetively. Table 37 contains an average for the winter va lues presented in Table 38 for the southern hemisphere. Furthermore da ta of Viebroek [363], Zillmann [364], Bunker [365], Winston et al. [366], Maykut [367], and Gavrilova [368] have been used (see also [369-372]).

For climate studies the deeisive question is whether or not the ice and snow cover of the earth remains constant in average or whether a trend can be observed. To answer this question very accurate measurements not only of the extent of ice and snow eoverage but also of the thiekness of the iee in polar regions are necessary.

The analysis which has been done by Lemke et al. [374] shows that the total ice cover in the Arctie deereased sinee the mid sixties by about 0.5· 106 km2

or ~ 5%. Also the antartic ice decreased by about 106 km2 or 10% between the early seventies and the la te seventies (Fig. 39). It eannot be assumed however,

2 E

I---If---lillf-.:..::m;o..hi=lla---.III-h"lfl,---j 0 .:

1973 1971. 1975 1976 1977 1978 1979

,----,,----,----,- -...,....----r--,---,--,--r--,----, 1.0

0.5

'" 0

E "'" '" Q

- 0.5

1966 1967 1968 1969 1970 1971 1972 1973 1974 1975 \976

Fig. 39. Variability of Arctic and Antartic ice cover. (After Lemke et al. [374])

o -2 -=

IV

-J

.J:o.

Tab

le 3

8. A

real

ly-a

vera

ged

obse

rved

hea

t and

rad

iati

on fl

uxes

. in

arct

ic r

egio

ns.

(Aft

er H

erm

an a

nd J

ohns

on [

362J

). A

ll va

lues

in

W m

-2

.

Sens

ible

hea

t L

aten

t hea

t S

olar

rad

iati

on

Net

lR a

t O

utgo

ing

Net

sol

ar

Net

flu

x flu

x ab

sorb

ed a

t su

rfac

e lo

ngw

ave

radi

atio

n su

rfac

e

1. Ja

nuar

y-F

ebru

ary

A.

Nor

ther

n H

emis

pher

e 1.

C

entr

al A

rcti

c -

1.6

-0.

4 0

-31

-1

50

to -

20

0

0 -1

50

to -

17

0

2.

Nor

weg

ian-

Bar

ents

sea

44

.1

48.1

o t

o 20

-2

0to

-3

5

-20

0 to

-2

25

O

to

50

-12

5 t

o -1

75

3.

L

abra

dor

Sea

35 t

o 70

50

to

200

5 to

20

-

2 to

-3

0

-17

5to

-2

00

o t

o 50

-1

25

to -

17

5

4.

Sib

eria

0

<1

5

5 to

100

-2

5 to

-4

0

-15

0to

-1

75

o t

o 50

-1

25

to -

17

5

B.

Sou

ther

n H

emis

pher

e 1.

A

ntar

ctic

con

tine

nt

8 0

100

-51

-1

75

to -

20

0

150

to 2

00

o to

20

2.

P

erip

hera

l oc

eans

(60

-70°

S)

3±

7

17

±7

15

4 -2

00

to -

22

5

17

5t0

22

5

40

to

80

3.

Per

iphe

ral

ocea

ns (5

0--6

0° S

) 1

8±

12

32

±3

17

2 -

3 to

-2

0

-20

0 tp

-2

25

20

0 to

250

40

to

80

4.

Wed

dell

Sea

-2

00

to -

22

5

150

to 2

00

11.

July

A

. N

orth

ern

Hem

isph

ere

1. C

entr

al A

rcti

c -

1.1

2.7

124

-29

-2

25

to -

25

0

150

to 2

00

Oto

21

2.

N

orw

egia

n-B

aren

ts s

ea

3.8

7.5

180

to 2

00

-15

to -

35

-2

25

to -

25

0

200

to 2

50

20 t

o 40

3.

L

abra

dor

Sea

-15

10

2

-5

to 2

0 10

0 to

200

-2

5 to

-3

5

-22

5 to

-2

50

25

0 to

300

2

0to

40

4.

S

iber

ia

o to

50

3010

70

218

to 2

40

-22

5 to

-2

50

25

0 to

300

40

to

80

B.

Sou

ther

n H

emis

pher

e 1.

A

ntar

ctic

con

tine

nt

-16

-

2 0

-20

-1

50

to -

17

5

0 -

150

to -

18

5

2.

Per

iphe

ral

ocea

ns (

60-7

0° S

) 20

to

40

-20

to -

30

-1

75

to -

20

0

4 -1

75

to -

20

0

3.

Per

iphe

ral o

cean

s (5

0--6

0° S

) 1

0±

12

24

±7

o t

o 20

-2

4

-20

0 to

-2

25

31

-1

50

to -

20

0

4.

Wed

dell

Sea

89

28

0

-10

to -

12

-1

75

to -

20

0

-17

5to

-2

00

p:: :.!..

. tl

j 2- "

Tab

le 3

9. T

he t

erra

in p

aram

eter

s, s

urfa

ee t

ype

and

siz

e of

live

dif

fere

nt l

atit

udin

al z

ones

bet

wee

n 0-

900

E in

Ant

aret

iea.

(A

fter

Wel

ler

[363

J)

Inte

rior

C

oast

al

Coa

stal

P

aek

iee

zone

eo

ntin

enta

l ab

lati

on

sea

iee

zone

zo

ne

zone

In

ner

Out

er

iee

eone

entr

atio

n 85

% >

iee

eon

eent

rati

on

~85%

~15%

Ter

rain

par

amet

ers

Alb

edo

85

69

65

75

54

Rou

ghne

ss l

engt

h (e

rn)

0.01

4 0.

26

0.01

5 0.

02

0.02

T

herm

al d

iffu

sivi

tyof

subs

trat

e (e

m2

s-1)

0.

008

0.01

3 0.

007

0.00

6 0.

0006

Su

rfae

e de

seri

ptio

n G

entl

e sl

opes

, S

teep

slo

pes,

F

ast i

ee, s

how

H

ighi

ee

Low

iee

Ope

n pe

rman

ent

blue

iee

surf

aee

dept

h <

5 e

m

eone

entr

atio

n,

eone

entr

atio

n,

oeea

n sn

oweo

ver

snow

dep

th

snow

dept

h >

5em

>

5em

Si

ze o

f zon

es

Ave

rage

wid

th (k

m)

2,50

0 20

15

60

0 50

0 P

eree

nt o

fto

tal a

rea

(%)

ofic

e eo

ver

47

0.5

0.5

26

26

betw

een

0-90

0 E

at

max

imum

iee

exte

nt

Net

rad

iati

on f

lux

at t

he s

urfa

ee (W

m-2

) -2

8 to

-1

2

-43

-5

6

-12

Se

nsib

le h

eat f

lux

(W m

-2)

12

to

25

39

5 -8

9

-15

2

-29

L

aten

t he

at f

lux

(W m

-2)

1

to

4 -

6 9

-28

-

37

-24

S

ub s

urfa

ee h

eat f

lux

(W m

-2)

2

to

3 10

42

65

~

~ i: 8- K

~ i S' Ir m ~ I f IV

-..I

VI

276 H.-J. Bolle

700 - Albedo 0.25 Q40

600

500

400

Sky 300 - -

N

200 , E 3: c:

100 x :J

0 >-Cl ~

<1.1 ------ --c: w -100

-200 Sensible heat

-300

-400 .... -----

- 500

True loeal t ime (h)

Fig. 40. Diurnal variation of energy balance components at different surfaces. Left: Model computations for two different albedos under identical solar irradiance. The effect of an increase in albedo from 0.25 to 0.40 is demonstrated. (After Seginer [377]). Right: Energy balance components on top of a low albedo tropical forest and below its canopy at the ground after Pinker et al. [251], and in the Tibetian highlands after Häckel et al. [247]


N

'E 3: c

x :J

C

" ""0

" Ir

"t E 3: c

X :J

>-~ ., c

W

1400

1200

1000

800

600

400

200

February

avg. net radiation (LW+SW)

avg. downward rad iation flux (LW+SW) at forest top

Forest Iloor

Or---------~~~------~~--~~--------

-100 U--t-+-+-~f-'1

o

1200 Chukhung

1000 6.4.1963

800

600

400

200

0

-200

-400

:~"""""'"'' .. , ...... / /'(/ -\. \ .... , ..........• ~Heat Ilux into soil

<:' ... ,/_ ...... _,.._. .~ Late~t heat \ .. ~........................... Sensible heat

0 True loeal time

277

278 H.-J. Bolle

that this is a continuing trend. The regression of the ice is not even uniform, there are longitudes where the ice proceeded significantly.

Though the annual variations of the ice and snow cover is a strong "signal", they are not the only surface changes which have to be taken into account if albedo - feedback processes are investigated. There are other and perhaps even irreversible processes in progress on the continents. Each change in land use affects the albedo of the surface - not to an extent as snow and ice cover does, but nevertheless in measurable dimensions. One transformation process is the conversion of tropical rainwoods first into arable land and probably afterwards into steppe. Another effect is the change of semi-arid agricultural land into desert due to dry periods and overgrazing as Ottermann [375] has pointed out. The albedo changes which occur presently in the Sahel zone are in the order of a few percent and their effect on climate is as difficult to assess as that of carbon dioxide. On the other hand also the irrigation of dry land changes the albedo, this time in the opposite direction.

The whole problem of land surface transformation has to be seen in an even wider aspect since here a number offeedback processes interact. A transformation into a dry surface is accompanied by a more intense generation of aerosols. Firstly aerosols are permanently produced due to buming ofwoods, and secondly soil particles can much easier be raised from the ground by the action of wind. This can lead to an aerosol albedo feedback. Furthermore the latent heat flux is reduced since the water supply can be expected to be reduced and the water level sinks. Thus the hydrological cycle is affected. There have been some assessment

0.38 14

0.34

6

0.30

0 0.26 "0 5

CIJ .0

<t 0.22 4

2,3

0.18 1 Day

0.14 Wet soil

0.1006 18 Loeal time (hl

Fig. 41. Albedo changes of smooth bare loam at Phoenix, Arizona, during a dry-out period. (After Idso [378])


studies [376], but to date it is difficult to judge the final effect of these transformations since also the magnitude of the albedo changes is not known to the necessaryaccuracy.

Seginer [377] has studied in a model computation what effect an increase ofthe soil albedo would have on the energy fluxes at the surface. The result is shown in Fig.40 in comparison with daily variations of the radiation budget in a tropical rain forest [251] and the Tibetian highlands [247]. Only few data are known about the processes which take place if a surface is drying out after rain or if the water level sinks during summer [378]. The drarnatic changes in albedo are demonstrated in Fig. 41. Two different effects can be seen here. Firstly the albedo change during the day in dependence of the solar zenith angle with the minimum at noon and an about 30% higher value at sunrise respectively sunset. Secondly the 100% albedo change (from about 15% to 30% atnoon) during one day if the water supply from below ceases. This occurs one day after rainfall in summer and a couple of days later in winter. Similar changes can be expected far large areas if forests (with very low albedo) are transformed to arable land and further to steppe and by erosion to desert.

The annual variability ofthe albedo has recently been estimated by Kukla and Robinson [379] separately for the oceans and land surfaces (Fig.42). Such data area necessary input for climate models since the albedo changes impose a forcing function with annual periodicity on such models which may yield different results than computations with annualy averaged values.

Monitoring of Climate Parameters

Monitoring Strategy

Monitaring of environmental parameters is done for different reasons. Firstly there is the problem to control at a local or regional scale the pollutants emitted from industrial sources which can be harmful for health, at least under extreme situations. Secondly there is the monitoring of the transport of pollutants over ranges of continental dimensions. Thirdly, and this becomes most essential in the case of climate relevant parameters, aglobai inventory of the distribution in space and time of the "critical" climate parameters is necessary.

The observation system to assess the variability and trends of climatically important parameters consists of different elements. (1) The monitoring ofpollutants on a local scale is in the responsibility of city or

state authorities. Rather den se networks exist in many communities with monitoring sites in the vicinity of main traffic lines, the centers of cities and in regions with concentration of industry.

(2) For larger areas certain sites have been selected which are expected to be representative for regions with respect to their air quality. These stations are coordinated by the World Meteorological Organization (WMO) jointly with the Uni ted Nations Environmental Programme (UNEP) in the WMO N etwork for Monitoring Background Air Pollution (Fig. 43).

a

80

---

1,..-

----

----

-1 __

_ ,

r------~

.I L--

1 I l __

, : r-J

I I I i !

I

5 L

_:

Ocean

r----

--l

I I

: L

____

_ --,

1

I 1

I 1

I I

I

r--

J "--

-1

I 1

I I

I 1

I l _

__

,

[----~

1

1

I r-

--J ~

I I L

__

_ _

I '/, BO

70

60

50

'0

30

20

\0

66

~

50

20-3

0 \0

~J"~

,,,~g

70

;ti!C

, '0

50

6~

80

80

90'L

-__

__

__

__

__

__

__

__

__

__

__

__

__

__

__

__

~

oS

oNIJ

Qn

FiI!'b

M

ar

Apr

M

o)'

Jun

Jul

Au~

S'P

O

el

Nov

o.c

l 84

I_~

_.

I I""

JI

----= :

::. ==

= =----

:::.:.

:::.~

l~ ll~

I r---

----

----

----

---

---

----

--

---

-15

L------

-I120

b

Fig.

42.

Var

iati

on o

fave

rage

zon

al m

onth

ly s

urfa

ce a

lbed

o in

the

cou

rse

of th

e ye

ar.

Fo

r a)

oce

ans

and

pack

ice

and

b) l

and.

Aft

er d

ata

give

n by

K

ukla

and

Rob

inso

n [3

79].

The

alb

edo

is r

epre

sent

ed i

n pe

rcen

t w

ith

10%

int

erva

ls.

The

num

bers

giv

en b

y K

ukla

an

d R

obin

son

are

mor

e de

tail

ed

IV

00

0 :Ii

~ l:l:I 2- (;"

30°1

1\

\

•

0°1-1

-----

-jl--

+--

f--t

----

-t--H

• o

()

30°1

11

R

egio

nal

air

pO

llutio

n st

ati

on

s CI

•

In o

per

atio

n

o

pLon

ned

Ilo

cate

dl c

r p

artl

yop

erat

lon

ol

o

pro

po

sed

(no

n lo

cQte

dl

Ba

selin

e a

ir p

ollu

tio

n s

tati

on

s

• In

op

erat

ion

I

I) 60

°1-

(J

plan

ned

(loc

ated

) or

par

tly o

pero

l,ona

l,-t-

--+-

-,.,--

-+--f'L

--t---+

--t---+

--I--

+--I--+

--I--

-+

--II---+

--1

--

o pr

opos

ed (

non

loca

tedl

1800

15

00

1200

9

0°

30° I)

2ö"'

S'W

75

"31'

5

0°

30°

• So

ulh

Pole

Fig

.43

. W

MO

Net

wor

k fo

r M

onit

orin

g B

ackg

roun

d A

ir P

ollu

tion

. (A

fter

[38

5bJ)

50°

90°

120°

15

00

1800

~ ~.

§ 8. m

~ ..., ~ ~

§ .;

o ::l

5' g- m ~ g o .; ir

@

rJ"J

'< g IV

00

.....

282 H.-J. Bolle

(3) In order to monitor the concentration of constituents in the most unperturbed air at the surface a small number of locations has been defined as Baseline Air Pollution Stations. These are as far away as possible from pollution sources such as industrial and agricultural centers, and are expected to reflect most clearly the global trends in concentration changes. They are part of the Background Air Pollution Network.

(4) The vertical distribution of air constituents is measured by special efforts on a research not an operational basis. Specially equipped research aircraft, stratospheric balloons, satellites and lidar stations contribute to these observations.

(5) Climate parameters are monitored by the operational meteorological network to which surface and aerological (radiosonde) stations and operational weather satellites as weIl as a few weather ships belong. These stations are maintained on a national basis but are coordinated by the World Meteorological Organization (WMO) within the World Weather Watch (WWW) and connected through the Global Teletype System (GTS).

The standard methods used for the chemical analysis of air and air pollution are described in great detail in the literature [381-386]. Less weIl known are the methods to establish the three dimensional picture of the distribution of air constituents and surface parameters which are of interest for climate research. These will therefore shortly be discussed here.

Baseline Stations

Because of its large geographical extension, its continous operation for already two decades, and its excellent maintenance the stations of the Uni ted States Geophysical Monitoring for Climate Change (GMCC) Program can be taken as exemplary for the global monitoring network. The U.S. network consists of the following baseline stations:

Mauna Loa Observatory, Mauna Loa, Hawaii Barrow Observatory, Point Barrow, Alaska Samoa Observatory, American Samoa South Pole Observatory, Antarctica

A number of special monitoring programs are carried out cooperatively with other, also non - U.S. institutions. In Table 40 a summary is given of the measuring programs of the Mauna Loa Observatory, which is the oldest one, and especially weIl known because of its CO2 record. CO2 is now monitored at a large number of stations (Table 41). Here the air is collected in paired 0.51 glass flasks and centrally analysed at the NOAA GMCC center in Boulder, Colorado, using an infrared CO2 analyzer [386]. A careful procedure is applied to the evaluation of the analysed sampies. Other gases which are sampled and centrallyanalyzed are CChF, CChF2, and N 20 [387].

Ozone is investigated optically by means of Dobson spectrometers [56, 389-390] using the "Umkehr" technique. In this method the zenith is observed at sunset when solar radiation penetrates the ozone layer at different optical pathes.


Table 40. Operational monitoring program at Mauna Loa in 1979. (After Herbert [342]). Additionally meteorological observations and cooperative programs are carried out

Monitoring program Instrument Sampling Remarks frequency

Gases Carbon dioxide URAS-2 infrared gas Continuous

analyzer Carbon dioxide Evacuated glass flasks Weekly Mountain and seacoast Surface ozone Electrochemical Continuous

concentration cell Surface ozone Dasibi ozone meter Continuous Totalozone Dobson spectrophotometer Discrete 3 meas. weekdays; 0 weekends Fluorocarbons Pressurized flasks Weekly

Aerosols Stratospheric Lidar Weekly 694.3 nm, 1 J

aerosols Condensation nuc1ei Pollak CN counter Discrete 5 meas. weekdays; 0 weekends Condensation nuc1ei G.E. CN counter Continuous Optical properties 4-wavelength nephelometer Continuous Wavelengths450, 550, 700, 850 nm

Solar radiation Global irradiance 4 Eppley pyranometers Continuous CutofTfilters at 280,390,530,

695 nm Ultraviolet irradiance Eppley ultraviolet Continuous Wavelength range 295 to 385 nm

pyranometer Direct beam Eppley pyrheliometer Continuous Wavelength range 280 to 3,000 nm

irradiance Direct beam Eppley pyrheliometer Discrete CutofT filters at 280, 530, 630,

irradiance with filters 695 nm Direct beam Eppley 13-channel Continuous

irradiance pyrheliometer

Precipitation chemistry

Activity ofrainwater pHmeter Discrete Rainwater collections at 6 sites Conductivity of Conductivity bridge Discrete

water Chemical Ion chromatograph Discrete

components

The ozone absorption in the ultraviolet part of the spectrum is observed spectroscopically. From the slope of the absorption curve at large solar zenith angles the ozone profile can be reconstructed. In addition in situ measurements of ozone are made from ballons. Here first a KI solution is oxidized by the ozone and the amount of released iodine is determined by an electrical conductivity measurement.

The monitoring program includes the analysis of precipitation and dry deposition at the surface. The pH-value of precipitation is determined monthly. For the Hawaii station, which is situated in the trade wind region, the annual mean pH values vary presently between 4.48 (3,400 m altitude) and 5.24 (0 m altitude). Also the monitoring of atmospheric turbidity and solar radiation is part of the observation program.

284 H.-J. Bolle

Table 41. COz flask sampling stations. (After Herbert [342J)

Station Country Lat.jlong. Elevation Agency (m)

SamoaObs. U.S. territory WI5'S,170034'W 30 (GMCC station) Amsterdam Is., Indian France 37°52'S, 77°32'E 150? Centre des Faibles

Ocean Radioactivities South Pole Obs. Antarctica 89°59'S, 24°48'W 2,810 (GMCC station) Ascension Is., South U.K. 7°55'S, W25'W 54 USAF /Pan American

Atlantic W orld Airways Azores (Terceira Is.), Portugal 38°45'N,27°05'W 30? U.S. Dept. Air Force

N. Atlantic 7th Weather Wing Cape Kumakahi, Hawaii U.S. W31'N, 154°49'W 3 (GMCCsite) Cold Bay, Alaska U.S. 55°12'N,162°43'W 25 NOAA,NWS Cosmos Peru 12°07'S, 75°20'W 4,600 Instituto Geofisico de Peru Easter Is., South Pacific Chile 27°06'S,I09°15'W 200 Ministro de Defensa

Nacional Direction Meteorologica de Chile

Falkland Is. U.K. 51°42'S, 57°52'W 51 British Meteorological Service

Galapagos (Chatham Is.), Equador 00054'S,89°37'W 6 Instituto National de South Pacific Meteorologia e

Hidrologia Guam (Marianas Is.), U.S. territory 13°26'N, 144°47'E

N. Pacific Key Biscayne, Pla. U.S. 25°40'N,80010'W 3 NOAA, Sea-Air Inter-

action Laboratory Mauna Loa, Hawaii U.S. W32'N,155°35'W 3,397 (GMCC station) Mauna Kea, Hawaii U.S. W50'N, 15S028'W 4,220 Institute for Astronomy,

Univ.ofHawaii Mould Bay, N.W.T. Canada 76° 14'N, 119°20'W 15 Dept. ofEnvironment,

Atmospheric Environment Service

Niwot Ridge, Colo. U.S. 40003'N, 105°38'W 3,749 INSTAAR, Univ. of Colorado

Ocean Station Charlie N. Atlantic 54°OQ'N,35°OQ'W 6 NOAA,NWS (U.S.)

Ocean Station M N.Atlantic 54°OQ'N,35°00'W 6 Norwegian Meteorologi-cal Institute

Palmer Station Antarctica 64°55'S, 64°OO'W 10? Desert Research Lab., (Anvers Is.) Univ.ofNevada

Point Barrow, Alaska U.S. 71°19'N,156°36'W 11 (GMCC station) Point Six Mountain, U.S. 47°02'N, 113°59'W 2,462 U.S. NWS

Mont. Seychelles (Mahe Is.), Seychelles 5°20'S,55°1O'E 3? Physical Science Lab.,

Indian Ocean New Mexico State Univ. St. Croix, Virgin U.S. Territory 17°45'N,64°45'W 3 Fairleigh Dickinson Univ.

Islands


Upper Atmosphere Monitoring

Measurements in the upper troposphere are mainly done by special aircraft missions. As an example, NASA in the U.S.A. has established the Aerosol Climatic Effects (ACE) program to obtain information on aerosols complementary to satellite and ground-truth programmes [391]. The measurements made from U-2 aircraft include sulfur compounds, N20, CO2, aerosol absorption coefficients and scattering phase functions, and aerosol size distributions.

There are two types of measurements made from stratospheric research balloons: spectrometric measurements of the attennuation of solar radiation by gases, and in situ probe sampling in evacuated steel flasks.

The flask sampling method has been described by Fabian et al. [392a] and by Fabian [392 b]. The latitudinal dependence of some gas constituents in the stratosphere have been investigated e.g. Goldan et al. [393a] and Farmer et al. [393 b]. For all radiatively important gases at least representative vertical profiles are known. The geographical distributions and temporal variations are still subject to numerous ongoing investigations.

The spectroscopic techniques which are applied from balloons [393c] as weIl as from satellites can be classified as transmission or occultation measurements and as (lirnb) emission measurements. The transmission technique makes use of the solar radiation which penetrates the atmospheric layers on top of the balloon and travels tangentially through the atmosphere at sunrise or sunset to a satellite instrument. The depth or equivalent width of the absorption lines [compare Eq. (75)] are determined from which the absorbing mass can be deduced [394-398]. The emission technique makes use of the long optical pathes near the earth horizon where the lines of rotation or rotation-vibration bands can be observed against space. The equivalent width of the lines of bands depends on the temperature structure of the atmosphere and the mass distribution. For an analysis of the mass distribution therefore also the vertical temperature profile must be kown.

By conducting scans either at different solar zenith angles or different air masses near the horizon it is possible to recover by application of mathematical inversion methods [399] the vertical mass distribution of the observed constituent. These measurements are either made in the infra red or in the microwave part of the spectrum.

Another optical technique which is applied for ozone monitoring from space is the backscatter method. Ultraviolet solar radiation is scattered in the atmosphere and the amount of radiance returning to space is analyzed by a satellite radiometer. The backscattered intensity at specific wavelengths where ozone absorbs depends on the height distribution of the absorbing molecules. By using different wavelengths with different penetration depths it is possible to deduce the vertical distribution of the absorbing molecules.

F or monitoring oftropospheric and stratospheric properties, primarely aerosol distributions, Lidar techniques are gaining in importance [400,25 d]. The analysis ofthe return signals from laser pulses sent into the atmosphere allow determination of aerosol layer stratifications and size distributions of aerosols. Only recently lidar systems have also been flown on board of aircraft. This allows to assess

286 H.-J. Bolle

within shart time the three dimensional structure of aerosol distributions in the troposphere [401, 402]. Because of its high power consumtion and technical constraints the lidar technique is not yet feasible far application in satellites but is a candidate for the Space-Shuttle.

On short ranges it is also possible to monitor gas pollutants by Lidar technique from the ground. In this case not the scattering efficiency of the constituent is used as indicator but either the absorption or the stimulated emission of gas specific lines. This technique is limited in range by the power of the laser since the return signals are very weak.

The chemical analytic methods for air constituent measurements are thus supplemented by physical sampling and analysis methods which develop their potential in cases where in si tu chemical measurements are difficult to apply. Combinations and intercomparisons of the different techniques are stilL very desirable in the future to improve further on the measuring accuracies and to enable a world wide standardization of the assessment of air pollutants and climate parameters.

List of Symbols

a) General Symbols

Symbol Definition Unit

a linear absorption coefficient rn-I

a = )./c(! thermal diffusity m2 s- 1

am mass absorption coefficient kg- 1 m2

aL line absorption coefficient m- 1 orkg- 1 m 2

ay absorption coefficient for a Voigt line rn-I or kg- 1 m2

A area m2

B self broadening coefficient C speed oflight m S-1

C specific heat capacity Jkg- 1 K- 1

CB concentration of substance B mol m- 3

cp specific heat capacity at constant pressure J kg- 1 K- 1

CV specific heat capacity at constant volume mol m- 3

C heat capacity Jm- 3 K- 1

e eccentricity e water vapor partial press ure Nm- 2

e elementary charge C E evaporation rate mmd- 1

E electrical field vector Vm- 1

f(v) frequency Hz 9 asymmetry factor G, G((8) geometry factor (e.g. cos ( or mean value over day etc.) h hour angle rad H = RT/(Mag) scale height m H magnetic field strength vector Am- 1

inclination angle degree n J source function Wm- 2

J, directional source function Wm- 2 sr- 1

k Boltzmann constant J K- 1

k von Karman constant


K eddy viscosity (I]/Q) m2 S-1

L latent heat J kg- 1

m mass kg MB molar mass of substance B kg mol- 1

M a molar mass of air kg mol- 1

M(() relative air mass at zenith angle ( n refractive index (n, real part, ni imaginary part) n number density m- 3

N total number density m- 3

N = n -I, refractive modulus N fractional cloud cover P pressure Pa P scattering phase function Pn residual polarization Pe equivalent pressure Pa p precipitation rate mmd- 1

p degree of polarization q specific humidity kg m- 3

Qeff efficiency factor Qe absorption efficiency factor Qs sca ttering efficiency factor

distance, radius m geometrical path length m

S line strength m- 2 or kg- 1 m S entropy J K- 1

S Poynting vector W t time T turbidity factor after Linke T (a bso I u te) tem pera ture K

T. equivalent radiation temperature K u optical path length m u zonal velocity m S-1

u. friction velocity m S-1

u(v) optical path at wavenumber v v velocity vector m S-1

v meridional velocity m S-1

W vertical velocity S-1

WB mass fraction of substance B kg/kg W equivalent width of a spectralline cm- 1

z height m Z geopotential height gpm ()( angle of incidence degree (0) ()( absorptance ()( specific volume kg- 1 m3

()( = ~ attenuation coefficient ()( solar azimuth angle degree C) ()( = 2nr/ A size parameter ()( halfwidth cm- 1

()(D Doppler half width cm- 1

()(L Lorenz half width cm- 1

()((v) absorptance at wavenumber v Ci absorption number

ß angle of reflection degree (0)

y =cJcv

y directional volume scattering function y specific conductivity S rn-I

YR volume scattering cross section for Rayleigh scattering m- 3

288

x A. A. A. 11 11 v v v*

12 12 ({J

<1>B(KB) w w Q Q

QO ( ,(v)

lapse rate (vertical) optical depth declination angle of earth axis perrnittivity (dielectric constant) ernittance viscosity, dynarnical angle between direction of incident electrornagnetic

wave and normal of irradiated area potential temperature =R/(Mh) geographicallongitude thermal conductivity wavelength rnagnetic permeability =cos ( ='1/12 viscosity, kinernatic wavenurnber slope of aerosol size distribution linear attenuation coefficient extinction coefficient due to Raleigh (rnolecular) scattering extinction coefficient due to Mie (aerosol) scattering linear scattering coefficient rnass scattering coefficient scattering cross section rellectance density geographicallatitude volurne fraction of substance B circular frequency (w= 2nf) albedo of single scattering angular velocity (of Earth rotation) solid angle unit solid angle zenith angle transrnittance at wavenurnber v transrnittance Ilux transrnittance rellection indicatrix

H.-J. Bolle

degree (0) Fm-I =rn- 3 kg- l s4 A2

K rn degree (0) Wrn-IK- I

rn,~rn

H rn - I = rn kg s - 2 A - I

Nsrn- 2

crn- I

rn-I rn-I rn-I rn-I kg- I rn2

rn- 2

kg rn- 3

degree n rn3/rn3

rad S-I

S-I

sr I sr degree C)

b) Symbols Used in the Description of the Energetics of the Atmosphere

Symbol

A

Q <1>=dQ

dt

M=d<1> dA

d<1> E= dA

Definition

Area Energy

Flux, Energy transport per unit time, power

Flux (area) density ( = exitance in case of a source of radiant energy)

Energy gain per unit area (= irradiance in case of radiant energy)

d2 <1> L= cos(} dAdQ Radiance, radiant energy Ilux per unit solid angle

through unit area oriented at an angle () to the direction of the Ilux

Unit

rn2

J, Ws

W


1= d<l> dQ

H= SEdt

Qp

B(T)

Intensity, radiant flux from point sources per unit solid angle

Radiant exposure, energy gain per unit area per specified time interval

Solar constant Irradiance of direct solar radiation

Albedo, ratio of radiation flux <1>, emerging from an area which is not emitting in the considered spectral range to flux <1>0 incident at the area

Planetary albedo at the top of the atmosphere Black body radiance of temperature T

c) Symbols Used in the Description ofEnergy Cycles

Symbol

G P C K D PE=gz IE=cvT LE =Lq SE =cpT

Iv l2

KE=T

U 2 +V2 KEh=-2-

w2

KEv=T

d) Indices

Symbol

tor -tor + * o

A adv

A a D d div

e gas GR g H

Definition

genera tion rate of energy per uni t area reservoir of available energy per unit area conversion rate into kinetic energy per unit area reservoir of kinetic energy dissipation rate per unit area potential energy per unit mass internal energy per unit mass latent energy per unit mass sensible energy per unit mass kinetic energy per unit mass

horizontal kinetic energy per unit mass

vertical kinetic energy per uni t mass

Definition

down up net value at a specified level ( = difference i ~ 1) sun infinite, also used for top of the atmosphere atmosphere advective: vertically integrated horizontal transport

of heat through a vertical area air, atmosphere aerosol Doppler line diffuse solar radiation Divergence of energy flux, sum of energy flux vectors over

all boundaries of a volume equivalent gas property ground (soil or seal diffuse and direct solar radiation = global radiation sensible plus latent heat

J m - 2 per defined time interval

Wm- 2

Wm- 2

Unit

Wm- 2

J m- 2

Wm- 2

J m- 2

Wm- 2

J kg- 1

J kg- 1

J kg- 1

J kg- 1

J kg- 1

289

290

h h

L LH LWorl m o o o

P PH r R R s SW SH v

v

horizontal high (ciouds) in, incomming low (ciouds) line Loren tz line latent heat longwave radiation, ..1.>2.5 Iilll medium high cIouds ocean reference level out planetary photosynthesis reflected Rayleigh ( = molecular) scattering total radiation (SW + LW, 0.2<). < 100 Iilll) surface shortwave radiation, 0.2< ..1. < 2.5 Iilll sensible heat vertical Voigt line

e) Mathematical Symbols

Symbol Name

square root of -1 -x x'

time average of quantity x

deviation from time average

[x] zonal mean

x deviation from zonal mean

<x) vertically integrated quantity

Frequently Used Numerical Va lues

1. Universal Constants

Name Symbol

Avogadro constant NA Molar volume, ideal gas Vm=RTo/Po Loschmidt constant no=N AiVm Molar gas constant R Boltzmann constant k=R/N A

Speed of light in vacuum c Permittivity of vacuum Go

Permeability of vacuum ~o

Stefan-Boitzmann constant (J

Planck constant h

H.-J. Bolle

Definition

.2.-l x dt ,1t t

x-x 1 2.

-2 S x dA n: 0

x-lxi 00

S x dp z=O

Numerical value Unit

6.0220. 1023 mo]-I

2.2414.10- 2 m3mol- 1

2.6867. 1025 m- 3

8.3144 J mol-I K- 1

1.3807 . 10 - 23 J K- 1

2.9979.108 m S-I 8.8542. 10- 12 Fm-I

m- 3kg- l s4 A2

4n: .10- 7 Hm- I m kg s-2A- 2

5.6703 . 10 - 8 Wm- 2K- 4

6.6262 . 10 - 34 Js


First radiation constant CI =2nhc2 3.7418.10- 16 Wm2 Second radiation constant C2 =hc/k 1.4388 . 10 - 2 mK Gravitational constant G 6.672 . 10- 11 N m2kg- 2

2. Geophysical and Astronomical Constants

Name Symbol Numerical value Unit

Astronomical Unit A.U. 149.6.106 km Excentricity of earth orbit e 0.01674 Mass ofsun mo 1.99. 1030 kg Radius of sun ro 696000 km Apparent radius of sun

1. July 15'45,42" 1. January 16'17,59"

Rotational period of sun Qo sideric 25.38 d synodic 27.27

Solar constant So 1370 Wm- 2 Mass of earth mE 5.974.1024 kg Large half axis of geoid rE, eq 6378 km

(equator) Small half axis of geoid rE,p 6357 km

(equator-pole)

Flattening of geoid f= rE,eq-rE, P

rE, eq 1 :297

Me,an orbital velocity of earth 29.8 km S-I

Angular velocity of earth Q=~ 7.2722.10- 5 rad S-I rotation 1 day (4.1667.10- 3 Os -I)

Mean siderial day (2n Q-I) (1+ 3~5) 86636.7

Rotational velocity at the v=rE' Q 464 m S-I earth surface (equator)

Acceleration due to gravity at sea level equator 90(0) 9.814 m S-2 pole 90(90) 9.832 m S-2 at 45°32'33" latitude 90(45) 9.80665 m S-2

Mass ofmoon mM 7.408. 1022 kg Mean distance earth-moon rE-M 384.420 km Exzentricity of lunar orbit 0.054908 Orbital lunar period

sideric 27d 07h 73' 12" synodic 29d 12h 44' 02,9"

3. Material Constants of Air and Water

Name Symbol Numerical value Unit

Apparent molar air mass M air 28.97 kg kmol- I

Gas constant for air Rair=RMai/ 287 J K-Ikg- I

Air density at STP !i.i, Qair,o 1.275 kg m- 3

(273.2 K, 10' Pa) Specific heat capacity of air Cp 1004 J K -I kg- I

at constant pressure (273 K)

292

Specific heat capacity of air Cv

Thermal conductivity of ). air at 273.2 K

Molar mass of water Gas constant of water

vapor Specific heat capacity of water

vapor at constant pressure at 273 K cwv• P

293 K Specific heat capacity of water

vapor at constant volume at 273 K Cwv• v

293 K Density of liquid water at

273.2 K Qw

277.2 K 293.2 K

Density of pure ice at Qi

273.2 K Specific heat capacity Cw

of pure water at 273.2 K (and const pressure)

Specific heat capacity Ci

of pure ice at 273.2 K

Latent heat of water vaporisation

2=2.5008 (273~5)

Latent heat of ice melting

~2.5 -0.002274 t t=T-273.16Cc)

2' = 334 (273 K) (:~330.7+2.03 t+O.OlO t2

717 2.40.10- 2

18.016 461.5

1847 1952

1386 1952

0.99984 . 103

0.99997 . 103

0.99820 . 103

0.917.103

4218

2106

Latent heat of ice 2"=2' +2~2.834-0.000149t 2.83 (273 K)

References

H.-J. Bolle

] K-1kg- 1 J m-1s-1K- 1

kg kmol- 1 J K-1kg- 1

kg m- 3

kg m- 3

kg m- 3

kg m- 3

1. Falk, G., Ruppel, W.: Energie and Entropie, Springer-Verlag, Berlin, Heidelberg, New York 1976 2. Fortak, H.: Entropy and Climate, in Bach, W., Pankrath J., Kellogg, W. (eds.): Man's Impact on

Climate, Development in Atm. Sciences, 10, Elsevier Pub\. Comp., Amsterdam, Oxford, New York 1979 •

3. Ekman, V.W.: Ark. Math., Astron. och Fysik 2,52 (1965) 4. Woods, J., 1981: The Memory of the Ocean, In: Berger, A., ed., 1981, Climatic Variations and

Variability: Facts and Theory. NATO Advanced Study Inst. Series, C, D. Reidel Pub\. Comp., Dordrecht, Holland, pp 63- 83

5. Maxwell, J.C.: A Treatise on Electricity and Magnetism, 2 Vols., Oxford 1873 6. Born, M., Wolf, E.: Principles ofOptics, Pergamon Press, London, New York, Paris, Los AngeIes

1978 7. Mie, G.: Beiträge zur Optik trüber Medien, speziell kolloidaler Metallösungen. Ann. Physik, Vier-

te Folge, 25, 377 (1908) 8. a Strutt, Hon. J.W. (Lord Rayleigh): Phi\. Mag. 4. SeI., 41, 107,274 (1871) 8. b Strutt, Hon. J.W. (Lord Rayleigh): Phi\. Mag. 4. SeI., 41, 447 (1871) 8. c Rayleigh, Lord: Phi\. Mag. 47, 375 (1899) 9. Liou, Kuo-Nan: J. Geophys. Res. 78, 1409 (1973)

10. Wendling, P., Wendling, R., Weickmann, H.K.: App\. Optics 18, 2663 (1979)


11. MeCartney, E.J.: Opties ofthe Atmosphere, John Wiley, New York, Londo, Sydney, Toronto 1976

12. Chandrasekhar, S.: Radiative Transfer, Dover Publ. Ine., New York 1980 13. Van de Hulst, H.C.: Multiple Light Scattering, Tables, Formu1as and Applications, Vo1.l and 2,

Aeademic Press, New York, London, Toronto, Sydney, San Franeisco 1980 14. Liou, Kuo-Nan: An Introduction to Atmospherie Radiation, Aeademic Press, New York 1980 15. Preisendorfer, R.W.: Hydrologie Optics, Vol.I-VI, U.S. Deptm. of Commerce, NOAA, ERL,

Honolulu, Hawaii 1976 16. Ishimaru, A.: Wave Propagation and Scattering in Random Media, Vol.l and 2, Academic Press,

New York, San Franciseo, London 1978 17. Hertz, H.: Wied. Ann. Physik Chemie, Neue Folge 36,1 (1889) 18. Valley, S.L. (ed.): Handbook of Geophysies and Space Environment, Air Force Cambridge Re-

search Laboratories, OAR, USAF 1965 19. Poynting, J.H.: Phi1. Trans. 11, 343 (1884) 20. Sommerfeld, A.: Elektrodynamik, Akad. Verlagsges. Geest u. Portig, K.-G., Leipzig 1949 21. Lorentz, H.A.: Ann Physik Chemie, Neue Folge 9, 641 (1880) 22. Lorenz, L.: Ann d. Phys., Neue Folge 11, 70 (1881) 23. Quenzel, H., Müller, H.: Optical Properties ofSingle Mie Partic1es: Diagram ofIntensity - Extinc

tion - Scattering - and Absorption Efficiencies. Münchener Universitäts-Schriften, Fachbereich Physik, Meteorolog. Inst., Wiss. Mitt. Nr.34, München (1978)

24. Junge, C.E.: Atmospheric Chemistry, in: Advances in Geophysies (Eds. Landsberg, H.E., Van Mieghem, J.) 1-108, Academic Press, New York 1958

25. a Foitzik, L., Spänkuch, D.: Z. Meteorol. 21, 136 (1969) 25. b Foitzik, L., Spänkuch, D., Unger, E.: PAGEOPH 69, 260 (1968) 25. c Foitzik, L., Scheithauser, G., Spänkuch, D.: Z. f. Meteorol. 21, 98 (1969) 25. d Russell, P.B., et al.: J. Atm. Sci., 38, 1279, 1296 (1981) 26. Deirmendjian, D.: Electromagnetic Scattering on spherical Polydisperions, Ameriean Elsevier

Pub1. Co., New York 1969 27. Pruppacher, H.R., Klett, J.D.: Microphysies of Clouds and Precipitation, D. Reidel Publ. Co.

Dordrecht, Boston, London 1978 28. Gerber, H.: Optical techniques for the measurement of light absorption by particulates, in: Par

ticulate Carbon. Atmospheric Life Cyc1e [G.T. Wolff, R.L. Klimisch (Eds.)] Plenum Press, N.Y., 1980

29. Lundgren, D.A. et a1.: Aerosol Measurement, Univ. Press of Florida, Gainesville, 1979 30. a Volz, F.E.: Appl. Opt. 11, 755 (1972) 30. b Volz, F.E.: J. Geophys. Res. 77, 1017 (1972) 31. Shettle, E.P., Fenn, R.W.: Models for the Aerosol of the Lower Atmosphere and the ElTects of

Humidity Variations on their Optical Properties, AFGL-TR -79-0214 Envir. Res .. Papers No. 676, AFGL - AFSC, USAF, Hanscom AFB, Mass. (1979)

32. Hänel, G.: Adv. Geophysics 19, 73, AcademiC"Press, New York 1976 33. Hänel, G., Bullrich, K.: Beitr. z. Phys. d. Atmosph. 51, 129 (1978) 34. Twomey, S.: Atmospheric Aerosols, Elsevier" Amsterdam, p.297, 1977 35. MeClatchey, R., Bolle, H.-J., Kondratyev, K.Ya.: A c10udless standard atmosphere for radiation

eomputations, IAMAP Radiation Commission, in preparation 36. Bolle, H.-J., Möller, F., Zdunkowski, W.: Investigation ofthe Infrared Emission Spectrum ofthe

Atmosphere and Earth, AF 61(052)-488, Techn. Rep. No. 2, Munieh, 1963 37. Herzberg, G.: Molecular Spectra and Molecu1ar structure, I. Spectra of Diatomic Molecules, II.

Infrared and Raman Spectra of Polyatomic Molecules, D. van Nostrand, Princeton, New Y ork 1950/1945

38. Penner, S.S.: Quantitative Molecular Spectroscopy and Gas Emissivities, Addison-Wesley, Read-ing, Mass., Pergamon Press, London, Paris 1959

39. Goody, R.M.: Atmospheric Radiation, Clarendon Press, Oxford, 1964 40. Unsöld, A.: Physik der Sternatmosphären, Springer-Verlag, Berlin 1955 41. Bolle, H.-J.: Beitr. z. Phys. d. Atmosph. 38, 41 (1965) 42. Bolle, H.-J.: Z. Geophys. 36, 1 (1970) 43. MeClatchey, R.A. et a1.: AFCRL Atmospheric Absorption Line Parameters Compilation,

AFCRL-TR-73-0096, Env. Res. Papers No. 434, AFSC-USAF, L.G. Hanseom Field, Mass. 1973

294 H.-J. Bolle

44. Rothman, L.S. et al.: Appl. Optics 20, 1323 (1981) 45. Chedin, A. et al.: La banque de donees "GEISA", description et logiciel d'utilisation, Laboratoire

de Meteorologie Dynamique du C.N.R.S., Note interna L.M.D. No. 108, Oct. 1980 46. Howard, J.N., Burch, D.E., Williams, D.: Infrared transmission of synthetic atmospheres. 11. Ab

sorption by carbon dioxide J.O.S.A. 46, 237-241 (1956); III. Absorption by water vapor J.O.S.A. 46, 242 (1956)

47. Burch, D.E., Gryvnak, D.A., Patty R.R.: J. Opt. Soc. Amer. 58, 335 (1968) 48. McCaa, D.J., Shaw, J.H.: J. Mol. Spectrosc. 25, 374 (1968) 49. Leupolt, A.: Infrared Physics 16, 523 (1976) 50. Roberts, R.E., Selby, J.E.A., Biberman, L.M.: Appl. Optics 15, 2085 (1976) 51. Deepak, A., Wilkerson, T.D., Ruhnke, L. (Eds.): Atmospheric Water Vapor, Academic Press,

New York, London, Toronto, Sydney, San Francisco 1980 52. a McClatchey, R.A. et al.: Optical Properties of the Atmosphere (3. Ed.), AFCRL-72-0497, Env.

Res. Paper No. 411, AFSC-USAF, Hanscom Field, Mass. 1972 52. b Carlon, H.R.: J. Atm. Scie 36, 832-837 (1979) 52. c Carlon, H.R.: Infrared Phys. 22, 43 (1982) 53. Banks, P.M., Kockarts, G.: Aeronomy, Part A and B, Academic Press, New York, London 1973 54. Nicolet, M.: Etude des reactions chimiques de I'ozone dans la stratosphere, lust. Royal Meteoro

logique de Belgique, Brussel 1980 55. CampelI, I.M.: Energy and the Atmosphere, John Wiley, London, New York, Sydney, Toronto

1977 56. Craig, R.A.: The Upper Atmosphere, Academic Press, New York, London 1965 57. Chamberlain, J.W.: Physics of the Aurora and air glow, Ißt. Geophys. Ser. 2, Academic Press,

New York, London 1961 58. Vassy, A.T., Vassy, E.: La luminescence nocturne, in: Encyclopedia of Physics (Ed. Flügge, S.)

Volume 49/5, 5 (1976) 59. McCormac, B.M., (Ed.): The Radiating Atmosphere, D. Reidel, Doordrecht 1971 60. McCormac, B.M. (Ed.): Physics and Chemistry of Upper Atmospheres, D. Reide1, Doordrecht

1973 61. Labs, D., Neckei, H.: Z. Astrophysik 55, 269 (1962) 62. Labs, D., Neckei, H.: Solar Physics 19, 3 (1971) 63. Thekaekara, M.P., Drummond, A.J.: Nature Phys. Sci. 229, 6 (1971) 64. Thekaekara, M.P.: Evaluating the light from the sun, Optical Spectra" March, 1972 65. Turner, C. Ed.: Solar Data Workshop, NSF-NOAA, Nat. Techn. Inf. Service, Washington, 1974 66. a White, O.R. (ed.): The Solar Output and its Variation, Colorado Ass. Univ. Press, Boulder 1977 66. bLondon, J., Fröhlich, C. (eds.): Solar Constant and the spectral distribution of solar irradiance.

Int. Rad. Comm. (IAMAP), Boulder, USA (1981) 67. NASA-SP-8005: Solar Electromagnetic Radiation, GSFC, Greenbelt, Maryland 1971 68. Wolfe, W.L., Zissis, G.J.: The Infrared Handbook, Env. Res. lust. Michigan, Office of Naval

Res., Washington D.C. 1978 69. Donelly, R.F., Pope, J.H.: NOAA Techn. Rep. ERL 276-SEL 25, USo Government Printing Of

fice, Washington, D.C. 1973 70. Friedman, H.: The sun's ionizing radiation, in: Physis of the Upper Atmosphere (Ratcliffe, J.R.

Ed.), pp 133, Academic Press, New York 1960 71. Kasten, F.: A new table and approximation formula for the relative optical air mass, CRREL

Techn. Rep. 136, U.S. Avrny Mat. Comm., Hanover, USA 1964 72. Linke, F., Baur, F.: Meteorologisches Taschenbuch, 11, 2.Aufl., 514, Leipzig 1970 73. Linke, F.: Die Sonnenstrahlung und ihre Schwächung in der Atmosphäre, in: Handbuch der Geo-

physik, Vol.8, Kap. 6 (Ed. Linke, F., Möller, F.), Gebr. Bornträger, Berlin 1942-1956 74. Quenzel, H.: Umschau 113,70 (1970) 75. a Yamamoto, G., Tanaka, M.: Appl. Optics 8, 447 (1969) 75. b King, M.D., et al.: J. Atm. Sci. 35, 2153 (1978) 76. a Wolfson, N., Joseph, J.H., Mekler, Y.: J. Appl. Met. 18, 543 (1978) 76. b Wolfson, N., Mekler, Y., Joseph, J.H.: J. Appl. Met. 18, 556 (1979) 77. Robinson, N. (Ed.): Solar Radiation, Elsevier, Amsterdam, London, New York, 1966 78. Schulze, R.: Strahlenklima der Erde, Steinkopff, Darmstadt 1970 79. Berlage, H.P.: Met. 45, 174 (1928)


80. Kondratyev, K.Ya.: The complete atmospheric energetics experiment, GARP Pub!. Series, No. 12, Geneva 1973

81. Leupolt, A.: Optik 24, 538, 567 (1966) 82. Bolle, H.-J., Leupolt, A., Möller, F.: Ausarbeitung eines Verfahrens zur Vorhersage thermischer

Strahlung, Forschungsauftrag Nr. 3119159, Beiheft, BMVtdg. (T II 3), München, 1963 83. Yates, H.W., Taylor, J.H.: J.O.S.A. 47, 223 (1957) 84. Bolle, H.-J.: Infrarotspektroskopie als Hilfsmittel und Gegenstand meteorologischer und planeta

rischer Forschung, BMWF-Forschungsber. W 67-17, München 1967 85. Völker, W.: Die Emission und Transmission der Atmosphäre im Wellenlängenbreich von 3-5 Ilm,

Diploma Thesis, Munich 1965 86. Bertram, F.-W., et al.: Ausarbeitung eines Verfahrens zur Vorhersage thermischer Strahlung, Un

tersuchungen zur Emission, Transmission und Reflexion infraroter Strahlung in der bodennahen Atmosphäre, Ber. V zum Forschungsauftrag T-489-1-203, BMVdtg, München 1967

87. Hane!, R.A., et al.: The NIMBUS infrared spectroscopy experiment, IRIS-D, Part I: Calibrated thermal emission spectra, NASA Prepr. X-622-71-272, G.S.F.C., Greenbelt 1971

88. Stephens, G.L.: J. Atm. Sei. 35, 2111 (1978) 89. Stephens, G.L., Paltridge, G.W., Platt, G.M.R.: J. Atm. Sei. 35, 2133 (1978) 90. Liou, Kuo-Nan, Wittman, G.D.: ibid. 36, 1261 (1979) 91. Korb, G.: Absorption von Sonnenstrahlung in Wolken, Wiss. Mitt. Nr. 6, Univ. München, Met.

Institut 1961 92. Dubrovina, A.S. (ed.): Aviation-Climatic Atlas-Guide ofthe USSR, Statistical Characteristics of

the Spatial and Microphysical Structure of Clouds, issue 3, 2 (1975) 93. McKee, T.B., Cox, S.K.: J. Atm. Sei. 31, 1885 (1974) 94. a Wendling, P.: ibid. 34, 642 (1977) 94. b Harshvardhan et a!.: J. Atm. Scie. 38, 2500--2513 (1982) 95. Feige!son, E.M.: Preliminary Radiation Mode! of the Cloudy Atmosphere, Acad. Sei. USSR,

Sect. Oceanology, Atmospheric Physics and Geography, Moscow 1977 96. Binenko, V.I., et a!.: Isv. Akad. Nauk SSSR, Fis. Atm. Okean 11 (1975) 97. Stephens, G.L.: J. Atm. Sei. 37, 435 (1980) 98. Wendling, P., Wendling, R., Weickmann, H.K.: App!. Optics 18, 2663 (1979) 99. Liou, Kuo-Nan: J. Geophys. Res. 78, 1409 (1973)

100. Yamamoto, G., Tanaka, M., Asano, S.: J. Atm. Sei. 27, 282 (1970) 101. Thompson, E.S.: Water Resources Res. 12,859 (1976) 102. Davies, J.A., Schertzer, W., Nunez, M.: Boundary Layer Met. 9, 33 (1975) 103. Kasten, F., Czeplak, G.: Solar and terrestrial radiation dependent on the amount and type of

eloud, Solar Energy 24,177-189 (1980) 104. Twomey, S.: The influence of aerosols on radiative properties of elouds, in (Pittock et al.; Eds.):

Climate Change and Variability, a Southem Perspective, Cambridge Univ. Press, p.281, 1976 105. Grassi, H.: Contr. Atm. Phys. 48, 199 (1975) 106. Arit, H., Bolle, H.-J.: Investigation ofthe infrared emission spectrum ofthe atmosphere and earth,

Part I: Angle dependent reflectivity of natural surfaces, 1-12 micron, Final Scient. Rep., AF 61 (052)-778, Munich 1968

107. Cox, c., Munk, W.: J. Opt. Soc. Am. 44, 838 (1954); J. Mar. Res. 13, 198 (1954) 108. Plass, G.N., Guinn, J.A.: App!. Opt. 14, 1924 (1975) 109. Plass, G.N., Guinn, J.A.: ibid. 15, 3161 (1976) 110. Preisendorfer, R.W.: 1. Quant. Spectrosc. Radiat. Transfer 11, 723 (1971) 111. Raschke, E.: Contrib. Atmos. Phys. 45, 2 (1972) 112. Köpke, P.: Bestimmung der atmosphärischen Trübung mittels geostationärer Satelliten, Münche-

ner Univers.-Schriften, Fachbereich Physik, Met. Inst., Wiss. Mitt. No.30, 1977 113. Lauscher, F.: Handb. Geophys. 7, 724 (1955) 114. Monahan, E.L.: J. Phys. Ocean 1, 139 (1971) 115. Payne, R.E.: J. Atm. Sei. 29, 959 (1972) 116. Anderson, E.R.: Energy-budget studies, Water-Ioss investigations: Vo!. I-Lake Hefner studies,

U.S. Geo!. Survey Circ., No. 229, 71-88 (1952) 117. Hollmann, R.: Studies on the albedo of the sea surfaces, Ph. D. Dissert., Rep. Met. Ocean. Geo

phys. Sei. Lab., TR-68-5 118. Grischenko, D.L.: Tr. GI. Geofiz. Observ., No. 80 (1959)

296 H.-J. Bolle

119. Kondratyev, K.Ya.: Radiation Characteristics ofthe Atmosphere and the Earth surface, NASA TT F-678, Amerind Publ. Co., New De1hi 1969

120. Burt, W.V.: J. Meteor. 11, 283 (1954) 121. Ter-Markariantz, N.E.: Tr. GI. Geofiz. Observ. No. 80, 1959 122. a Ty1er, J .E., Smith, R.C.: Measurements of spectra1 irradiance underwater, Gordon and Breach,

New York, London, Paris 1970 122. b. Smith, R.C., Baker, K.: Appl. Optics 20, 177 (1981) 123. a Kislovskii, L.D.: Optics and Spectroscopy 7, 201 (1960) 123. b Irvine, W.M., Pollack, J.B.: Icarus 8, 324 (1968) 124. Bell, E.E.: An atlas of reflectivities of some common types of materials, Int. Engineering Rep.,

659-6. Wright Air Dev. Center Contract AF 33 (616)-3312, 1961 125. Köpke, P.: Contr. Atm. Phys. 53, 442 (1980) 126. Downing, H.O., Williams, D.: J. Geophys. Res. 80, 1656 (1975) 127. Coulson, K.L., Bouricius, G.M.B., Gray, E.O.: Effects of surface reflection on radiation emerging

from the top of a planetary atmosphere, Rep. 1, May, 1965; Rep. 2, Sept. 1965 128. Coulson, K.L., KinseIl, L.: Appl. Optics 5, No. 6, June (1966) 129. Bolle, H.-J., Beffert, R.: Experimente für meteorologische Satelliten oder eine We1traumstation.

Forschungsbericht BMBW W 70-70, München 1970 130. Hovis, W.A.: Infrared reflectivity of some common minerals. Appl. Optics 5, 245-248, 815-818

(1966) 131. Hoffer, R.M., Johannsen, c.J., in: Remote Sensing in Ecology, Univers. Georgia Press, Athens,

GA 1969 132. Bowers, S.A., Hanks, R.J.: Soil Science 100, 1301 (1965) 133. Kriebei, K.Th.: Reflection Properties ofVegetated Surfaces: Tables ofMeasured Biconical Reflec

tance Factors, Münchener Univ. Schriften, Fachbereich Physik, Meteorol. Inst., Wiss. Mitt. Nr.29, München 1977

134. Romanova, M.A.: Air Survey of Sand Deposite by Spectral Luminance, Consultats Bureau, New York 1964

135. quoted by: Bolle, H.-J.: Nutzung von Fernerkundungsdaten für die Klimaforschung, in: Mondre, E., Pollanschutz, J.: Österreich. Symp. Fernerkundung, Mitt. Forstl. Bundes-Versuchsanstalt 135, Wien 1981

136. Otterman, J., Fraser, R.S. Remote Sensing of the Environment 5,247 (1976) 137. Target Signatures Study Interins Report, Volume V: Catalog of Spectral Reflectance Data, The

University of Michigan. Ann Arbor, MI, Rep. No. 5698-22-T(V) (1964) 138. Büttner, K.J.K., Kern, C.D.: J. Geoph. Res. 70, 1329 (1965) 139. Kondratyev, K.Ya.: The albedo and the angular characteristics ofreflectance for the surface and

c10uds (in Russian), Gidrometeoizdat Leningrad 1981 140. Wiscombe, W.J., Warren, S.G.: J. Atm. Sei. 37, 2712 (1980); ibid: 37, 2734 (1980) 141. Kuhn, M., Siogas, L.: Antarctic J.U.S. 13, 178 (1978) 142. Wagner, H.-P.: Arch. Met. Geoph. Biokl. Ser. B 27, 297 (1979); ibid. 28, 41 (1980) 143. Sauberer, F., Dirmhirn, 1.: Geogr. Ann. 34, 261 (1952) 144. Scheibner, F., Mahringer, W.: Arch. Met. Geophys. u. Bioklim., Ser. B, 16, 174 (1968) 145. Bolle, H.-J.: Infr. Physics 5, 115 (1965) 146. Möller, F.: Strahlung in der unteren Atmosphäre, in: Handb. Physik, Vo148 (Ed. Flügge, S.),

Springer-Verlag, Heidelberg, Berlin 1957 147. Kondratyev, K.Ya.: Radiative heat exchange in the atmosphere, Pergamon Press, Oxford, Lon

don, Edinburgh, New York, Paris, Frankfurt 1965 148. Paltridge, G.W., Platt, C.M.R.: Radiative Processes in Meteorology and Climatology. Dev. in

Atm. Sei. 5, Elsevier 1976 149. Lenoble, J., Ed.: Standard Procedures to Compute Atmospheric Radiative Transfer in a Scatter

ing Atmosphere. Int. Radiation Comm., IAMAP, Boulder 1977 150. Fouquart, Y., Irvine, W.M., Lenoble, J (Eds.): Standard Procedures to Compute Atmospheric

Radiative Transfer in a Scattering Atmosphere, Vol. 11. Part A: Review of Methods for Horizontally Inhomogenous Atmospheres and Spherical Atmospheres. Part B: Problems of Scattering with Gaseous Absorption. Radiation Comm., IAMAP, Boulder 1980

151. Plass, G.N., Kattawar, G.W.: J. Atm. Sei. 28, 1187 (1971), J. Phys. Ocean. 2, 139 (1972) 152. Kattawar, G.W., Plass, G.N., Guinn, J.A.: J. Phys. Ocean. 3, 353 (1975)


153. Collins, D.G., et al.: App. Optics 11, 2684 (1972) 154. Marchuk, G.I., et al.: Solution of Direct and some inverse Problems of Atmospheric Optics by

the Monte Carlo Method, Novosibirsk 1968 155. Marchuk, G.I., Mikhailov, G.A.: Izv. Akad. Nauk. SSSR, Seria, Fizika atm. i. okeana, 3, 258

(1967) [Eng. trans. Izv. Ac. Sc. USSR, Atm. Oceanic Phys. 3, 167 (1967)] 156. Canosa, J., Penafiel, H.R.: J. Quant. Spectro. Rad. Trans. 13,21 (1973) 157. Deuze, J.L., Devaux, C., Herman M: Nouv. Rev. d'Optique 4, 307 (1973) 158. a Nagel, M.R., et al.: Tables related to the illumination, color and contrast in naturally illumi

nated objects, Academic Press, New York 1978 158. b Zdunkowski, W.G., Korb, G.: Contributions Atm. Phys. 47, 129 (1977) 159. Carlson, B.: The Numerical Theory ofNeutron Transport, in: "Methods ofComputational Phys-

ics", New York, London 1963 160. Chandrasekhar, S.: Radiative Transfer, Clarendon Press, Oxford 1950 161. Kuznetsov, E.S.: Izv. Akad. Nauk SSSR ser Geo. 71,91 (1951) 162. Liou, K.N.: J. Atm. Sci. 30, 1303 (1973) 163. Braslau, N., Dave, J.V.: J. App. Meteor. 12, 601 (1973) 164. Danzer, K.M., Bullrich, K.: Beitr. zur Phys. Atm. 41, 143 (1968) 165. Eschelbach, G.: J. Quant. Spectro. Rad. Transfer 11, 757 (1971) Ann. Geophys. 29, 329 (1973) 166. Herman, B.M., Browning, S.R., Curran, R.J.: J. Atm. Sci. 28, 419 (1971) 167. de Bary, E.: App. Optics 3, 1293 (1964) 168. Dave, J.V., Furukawa, P.: J. Opt. Soc. Amer. 56, 394 (1966) 169. Uesugi, A., Irvine, W.M.: Astroph. J. 159, 127 (1970),161,243 (1970) 170. Germogenova, T.A.: Soviet Math. Dokl. 9, 855 (1968) 171. Van de Hulst, H.C.: Astron. Astroph. 9, 374 (1970) 172. Grant, I.P., Hunt, G.E.: Proc. Roy. Soc. A313, 183, 199 (1969) 173. Takashima, T.: Astroph. Space Scie. 23, 201 (1973) 174. Kattawar, G.W.: J. Quant. Spectro. Rad. Trans. 13, 145 (1973) 175. Plass, G.N., Kattawar, G.W., Catchings, F.E.: App. Optics 12, 314 (1973) 176. Preisendorfer, R.: Radiative Transfer in Discrete Spaces, Pergamon Press 1965 177. Hansen, J.E., Hovenier, J.W.: J. Quant. Spectro. Rad. Trans. 11, 809 (1971) 178. Lacis, A.A., Hansen, J.E.: J. Atm. Sci. 31, 118 (1974) 179. Takashima, T.: J. Quant. Spectro. Rad. Trans. 13, 1229 (1973) 180. Tanaka, M.: J. Meteor. Soc. Japan. 49, 296, 321, 333 (1971) 181. Hansen, J.: J. Atm. Sci. 28, 120 (1971) 182. Danielson, R.E. et al.: J. Atm. Scie 26, 1078 (1969) 183. Twomey, S., Jacobowitz, H., Howell, H.B.: J. Atm. Sci. 23, 289 (1966) 184. Ivanov, V.V.: Astron. Zh. 52, 217 (1975) 185. Lenoble, J.: Methode des Principles d'Invariance pour une Couche Semi-Infinie, in: Comparaison

des Methodes de Resolution de I'Equation de Transfer: I. Sans Polarisation, Rapport Universite des Sciences et Techniques de Lilie. Dec. 1972

186. Adams, C.N., Kattawar, G.W.: J. Quant. Spectro. Rad. Trans. /0, 341 (1970) 187. Bellman, R.: Invariant Imbedding and Computational Methods in Radiative Transfer, in Trans

port Theory, Vol.I, SIAM-AMS Proceed. (ed. by Bellman, R .. , Birkhoff, G., Abu-Shumays, 1.), Providence. Rhode Island 1969

188. Bellman, R., Kalaba, R., Prestrud, M.C.: Invariant Imbeeding and Radiative Transfer in Slabs of Finite Thickness, Amer. Elsevier, New York 1963

189. Preisendorfer, R.W.: Radiative Transfer on Discrete Spaces, Pergamon Press 1965 190. Ueno, S., Mukai, S., Wang, A.P.: Invariant Imbedding and Chandrasekhar's Planetary Problem

ofPolarized Light, in Planets, Stars, Nebulae Studied with Photopolarimetry (ed. T. Gehreis), The Univ. of Arizona Press, p. 582 1973

191. Whitney, C.K.: J .. Quant. Spectrosc. Radiat. Transfer 14, 591 (1974) 192. Van de Hulst, H.C., Grossman, K.: Multiple Light Scattering in Planetary Atmospheres, in The

Atmospheres of Mars and Venus (ed. Brandt, J.c., McElroy, M.B.), Gordon and Breach, New York 1968

193. Kawata, Y., Irvine, W.M.: Astroph. J. 160,787 (1970); Shettle, E.P., Weinman, J.A.: J. Atm. Sci. 27, 1048 (1970)

194. a Irvine, W.M.: Astroph. J. 152, 823 (1968)

298 H.-J. Bolle

194. b Wiscombe, W.J.: J. Atm. Scie. 34, 1408 (1977) 195. Korb, C., Michalowsky, J., Möller, F.: Beitr. Phys. Atm. 30, 63 (1957) 196. Liou, K.N.: JGR, 78, 1409 (1973) 197. Schuster, A: Astroph. J. 21, 1 (1905) 198. Zege, E.P.: Two-Stream Approximation in Transfer Radiation Theory, Preprint Minsk-Inst.

Phys. Acad. Sci., 1971 199. Bigourd" D., Devaux, C., Herman, M.: Albedos plan et spherique-Extension de la Methode de

Wang (Noyau Exponentie1), Rapport Univers. Sciences et Techniques de Lilie, Nov. 1973 200. Wang, L.: Astroph. J. 174, 671 (1972) 201. Deuze, J.L., Devaux, C., Herman, M.: Nouv. Rev. d'Optique 4, 307 (1973) 202. Fymat, AL., Abhyankar, K.D.: Astroph. J. 158, 315, 325 (1969); 159, 1009, 1019 (1970) 203. Sobolev, V.V.: Light Scattering in a Planetary Atmosphere, Izd. Nauka, Moscow. (Eng. trans.

Pergamon Press 1975), Chapters V and VI, 1972 204. Bakan, S., Quenzel, H.: Beitr. Physik d. Atmosphäre 51, 15 (1978) 205. Kneizys, F.x., et al.: Atmospheric Transmittance/Radiance: Computer Code LOWTRAN 5;

AFGL-TR-0067, Environmental Res. Papers, No. 697, Air Force Cambridge Geophysics Lab., AFSC-USaf, Hanscom AFB, Mass. 1980

206. Chedin, A, Scott, N.A.: Quantitative analysis of radiometric measurements from satellites, new computationally fast line-by-line transmittance and radiance model, Proc. Int. Radiatoin Symposium, Ft. Collins, Colorado, p 199, 1980

207. Curtis, AR.: Quart. J. Royal Meteorolog. Soc. 78,638 (1952) 208. Godson, W.L.: ibid. 79,367 (1953) 209. Weinreb, M.P., Neuendorfer, A.C.: J. Atmosph. Sci. 30, 662 (1973) 210. Yamamoto, G., Aida, M.: J. Quant. Spectr. Radiat. Transfer 10, 99 (1970) 211. Elsasser, W.M., Culbertson, M.F.: Atmospheric Radiation Tables, Univers. California, La Jolla,

AFCRL-TR-60-236, 1960 212. Rodgers, C.D., Walshaw, C.D.: Quart. J. Royal Meterolog. Soc. 92, 67 (1966) 213. Paltridge, G.W.: ibid. 101,475 (1975) 214. Oort, AH., Vonder Haar, T.H.: J. Phys. Ocean. 6, 781 (1976) 215. Vonder Haar, T.H., Suomi, V.A: J. Atmos. Sei. 28, 305 (1971) 216. Raschke, E., et al.: ibid. 30, 341 (1973) 217. Vonder Haar, T.H. and V.E. Suomi: ibid. 28, 305 (1971) 218. London, J., Sasamori, T.: Radiative budget of the atmosphere, Space Research IX, Akademie

Verlag, Berlin, p. 639 (1971) 219. Sasamori, T., London, J., Hoyt, D.V.: Radiation budget of the southern hemisphere, Meteor.

Monographs 13, 9 (1972) 220. Kendall, J.M.: Proc. Symp. on Solar Radiation, Smithsonian Inst., Washington D.C. 1973 221. Murcray, D.G.: Balloon Borne Measurement ofthe Solar Constant, Rep. No. AFRCL-69-0070,

Univers. Denver, Col. 1969 222. a Kondratyev, K.Ya., Nikolsky, G.A.: Quart. J.R. Met. Soc. 96, 509 (1970) 222. b Kondratyev, K.Ya. et a1.: The solar constant from data of ballon investigations in the USSR

and the USA, Space Research XI, 695, Akademie-Verlag, Berlin 1971 223. Willson, R.C.: Appl. Optics 12, 810 (1973) 224. Brusa, R.W., Fröhlich, C.: Monitoring the Solar Constant from high altitude balloons, Proc. Int.

Rad. Symp. IAMAP, Colorado State Univers., Ft. Collins, Colorado, USA, 416, 1980 225. Willson, R.C., Duncan, C.H., Geist, J.: Science 207,177 (1980) 226. Plamondon, J.A.: JPL Space Program, Summary 3, 162 (1969) 227. Smith, W.L., et al.: Appl. Opt. 16, 306 (1977) 228. Hickey, J.R., et al.: Seience 208, 281 (1980) 229. Jacobowitz, H., et al.: J. Atm. Sci. 36, 501 (1970) 230. Vonder Haar, T.H., et al.: Measurements ofthe earth radiation budget from satellites during the

first GARP global experiment, in: Tänczer, T., Götz, G., Major, G.: First FGGE Results from Satellites, Adv. in Space Res., Vol.l, No. 4, p. 285, Pergamon Press, Oxford 1981

231. Fröhlich, C.: Contemporary measures of the solar constant, in: The Solar Output and its Variation (White, O.R., Ed.), 93-109, Colorado, Ass. Univ. Press, Boulder 1977

232. a Brusa, R.W., Fröhlich, C.: The solar constant: recent results, Int. Pyrh. Comp. 1980; Arb.-Ber. d. Schweiz. Met. Anstalt Zürich, Nr. 94, Davos und Zürich (1981), 9; Jahresber. Phys.-Met. Obs. and World Radiation Center Davos 1981


232. b Crommelynck, D.: Approche objective d'une valeur de la constante solaire validee par determination indirecte de la constante de Stefan-Boltzmann. Inst. Roy. Met. de Belgique, Pub!. Ser.A., No. 91,95-106, 1975

232. c Fröhlich, c., Brusa, R.W.: Solar Physics 74,209 (1981) 232. d Willson, R.C., et a!.: Science 211,700 (1981) 233. Hinteregger, H.E.: AE-E experiments of irradiance monitoring for 1200-1850 A, NOAA Work-

shop on Solar UV-Monitoring, Boulder, Co!. 1980 234. Braslau, N., Dave, J.V.: J. App!. Met. 12, 601 (1973) 235. Braslau, N., Dave, J.V.: ibid. 12, 616 (1973) 236. Raschke, E., Bandeen, W.R.: ibid. 9, 215 (1970) 237. Campbell, G.G., Vonder Haar, T.H.: Climatology of Radiation Budget Measurements from Sa

tellites, Atm. Science Paper No. 323, Dptm. Atm. Sei., Colorado State Univers., Ft. Collins, Colorado 1980

238. Ellis, J.S., et a!.: J. Geophys. Res. 83 C, 1958 (1978) 239. Barrett, E.C.: Climatology from Satellites, Methuen, London 1974 240. Bolle, H.-J.: Data for radiation climatology (Guyenne, T.D., Levy, G. Eds.), Proc. 2nd Course

on Satellite Meteorology ofthe Mediteranean. ESA SP-159, Paris 1981 241. Budyko, M.l.: Atlas teplovogo balansa zemnogo shara, Mezhduvedomstvennyi Geofizicheskii

Komitet pri Prizidium, Akademiia Nauk SSSR, Glavnaia Geofizicheskaia Observatoriia imennii A.E. Voeikova, Reszultaty

242. Budyko, M.l., et a!.: Sov. Geog. 3, (1962) 243. Perry, A.H., Walker, J.M.:: The Ocean-Atmosphere System, Longman, London, New York 1977 244. Vowinkel, E., Orvig, S.: The climate of the north polar basin, in: World Survey of Climatology,

Vo!. 14, Elsevier Pub!., Amsterdam, London, New York 1970 245. Marskunova, M.S.: Principal characteristics ofthe radiation balance ofthe underlying surface and

ofthe atmosphere in the Arctic (in Russian), Trudy AAHUU T 229, Leningrad 1963. See: Fleteher, J.O.: The heat budget ofthe Arctic Basin and its relation to climate, Rep. R-444-PR, The Rand Corp., Santa Monica 1965

246. Schlatter, T.W.: J. App!. Met. 11, 1048 (1972) 247. Häckel, H., Häckl., K., Kraus, H.: Khumbu himal 7, 71 (1970) 248. Bowen, l.S.: Phys. Rev. 27, 779 (1926) 249. Baumgartner, A.: Climate Variability and Forestry, Proc. World Climate Conf., WMO No. 537,

p. 581, Geneva 1979 250. MitcheII, J.M., Jr.: J. App!. Met. 10, 703 (1971) 251. Pinker, R.T., Thompson, O.E., Eck, T.E.: ibid. 19, 1341 (1980) 252. Jarvis, P.G., James, G.B., Landsberg, J.J.: Coniferous forest, Vegetation and the Atmosphere,

Vo!. 2, Casc Studies [Monteith, J.L. (Ed.)], Academic Press, 171, 1976 253. Agee, E.M., Howley, R.P.: J. App!. Met 16, 443 (1977) 254. Holland, J.Z.: J. Phys. Oceanogr. 2, 476 (1972) 255. Wijk, W.R. van, Vries, D.A. de: Periodic temperature variations in a homogenous soil, in: Physics

ofPlant Environment, p. 171, North-Holland, Amsterdam 1963 256. Seilers, W.D.: Physical Climatology, The Univers. Chicago Press, Chicago, London 1965 257. Swinbank, W.C.: J. Meteoro!. 8, 135 (1951) 258. Hoffmann, U.: Probleme des Stadtklimas von Stuttgart, in: Stadtklima, [E. Franke (ed)], p. 65,

1977 259. Wilson, W.: J. Geophys. Res. 65, 3377 (1960) 260. Wu, J.: ibid. 74,444 (1969) 261. Wucknitz, J.: Meteor.-Forsch.-Ergebnisse B, 11, 25 (1976), Berlin, Stuttgart 262. Amorocho, J., DeVries, J.J.: J. Geophys. Res. 85C, 433 (1980) 263. Francey, R.J., Garratt, J.R.: Boundary Layer Meteoro!. 14, 153 (1978) 264. Bill, R.G., et a!.: J. Geophys. Res. 85, 507 (1980) 265. Penman, H.L.: Proc. Roy. Soc. A 193, 120 (1948) 266. a Thornthwaite, C.W.: Geogr. Rev. 38, 55 (1948) 266. b Trewartha, G.T.: Elements of Physical Geography, McGraw Hili, 1957 267. Anderson R.J., Smith, S.D.: J. Geophys. Res. 86 C, 449 (1981) 268. Lemon, E.R.: Agron. J. 52, 697 (1960) 269. Denmead, O.T.: Agricu1t. Meteoro!. 6, 357 (1969) 270. Budyko, M.l.: Climate and Life, Academic Press, New Y ork, London 1974

300 H.-J. Bolle

271. Winston, J.S.: Tellus 7, 481 (1955) 272. Starr, V.P.: Applications of energy principles to the general circulation, Compendium of Me

teorology, Boston, Amer. Meteor. Soc., p. 568, 1951 273. Newell, R.E., et al.: The energy balance ofthe global atmosphere, The Global Circulation ofthe

Atmosphere, Roy. Meteor. Soc. (London) p. 42, 1969 274. Oort, A.H.: J. Atm. Sci. 28, 325 (1971) 275. Wiin-Nielsen, A., Brown, JA, Drake, M.: Tellus 16, 168 (1964) 276. Lorenz, E.N.: Tellus 7, 157 (1955) 277. a London, J., Ohring, G., Ruff, I.: Radiative properties of the stratosphere, Final Rep. Contr.

AF 19 (604)-1285, Cambridge, Mass. 1956 277. bLondon, J.: Radiative Heat Sources and Sinks in the Stratosphere and Mesosphere. Proc.

NATO Adv. Study Inst. on Atm. Ozone: Hs Variation and Human Influences, Rept. No. FAAEE-80-20, 1980

278. Thomas, L.: Some outstanding problems in the neutral and ionized atmosphere between 60 and 150 km altitude, in: Atmospheric Physics from Spacelab, [Burger, J.J., A. Pedersen and B. Battrick (ed.)] D. Reidel Pub!., Dordrecht, p. 408, 1976

279. Markov, M.N.: App!. Optics 8, 887 (1969) 280. Kutepov, A.A., Shved, G.M.: Radiative Transfer in the 15 J.11l1 CO2 band with the non-LTE in

the Earth's atmosphere, Akademia Nauk 14, 1,28 (1978) 281. Grossmann, K.U., Offermann, D.: Nature 276, 594 (1978) 282. Mayr, H.G., Harris, 1., Spencer, N.W.: Rev. Geoph. Space Phys. 16, 539 (1978) 283. Hadley, G.: Concerning the Cause ofthe General Trade-Winds, Phi!. Trans. Roy. Soc. (London)

p. 39 (1735--36). Reprinted in: The Mechanics ofthe Earth's Atmosphere (A Collection ofTranslations by Cleveland Abbe) 3rd Collection, Smithson. misc. Coll., Vo!. 51, NO.4 (1910)

284. Palmen, E.H., Riehl, H., Vuorela, L.A.: J. Meteor. 15, 271 (1958) 285. Petrossiants, M.A., et a!.: The air circulation in the tropical troposphere along the meridian of

23°30'W, in: Preliminary Scientific Results of the GARP Atlantic tropical experiment, Vo!. 1. GATE Rep. No. 14, ICSU-WMO, Geneve 1975

286. Dietrich, G., et al.: Allgemeine Meereskunde, Gebr. Bornträger, Berlin, Stuttgart 1975 287. a Busalacchi, A.J., O'Brien, J.J.: J. Phys. Ocean. 10, 1929 (1980) 287. b Wyrtki, K.: J. Phys. Oceanogr. 5, 450 (1975) 287. c Walker, G.T.: Mem. Indi. Met. Dptm. 24, 275 (1924) 287. dRamage, C.S., Hori, A.M.: Monthly Wea. Rev. 109, 1827 (1981) 287. e Swanson, G.S., Treuberth, K.E.: Monthly Wea. Rev. 109, 1879, 1890 (1981) 287. f Wyrtki, K.: J. Phys.Oceanogr. 9, 1223 (1979) 288. Hastenrath, S.: J. Phys. Ocean 10, 159 (1980) 289. Bryden, H.L., Hall, M.M.: Science 207, 884 (1979) 290. Roemmich, D.: J. Phys. Ocean. 10, 1972 (1980) 291. Stomme1, H.: Proc. Nat. Acad. Sei., USA 77(5), 2377-2381, 1980 292. a Washington, W.M., et a!.: J. Phys. Ocean. 10, 1887 (1980) 292. b Gates, W.L. et al.: J.G.R. 86, C 7,6385-6393 (1981) 293. Sverdrup, H.U.: Evaporation from the oceans. In: Compendium of Meteorology (Ed. Malone,

T. F.), Amer. Meteor. Soc., p. 1071, 1951 294. Rakipova, L.R.: Izv. Atmos. Oceanic Phys. 2, 983 (1966) 295. Peixoto, J.P., Salstein, D.A., Rosen, R.D.: J. Geophys. Res. 86, 1255 (1981) 296. Campbell, G.G., Von der Haar, T.H.: Latitude average radiation budget over land and ocean

from observations and some implications for energy transport and climate modelling. Monthly Wea. Rev. in press 1981

297. Milankovitch, M.: Mathematische Klimalehre und astronomische Theorie der Klimaschwankun-gen, in: Handbuch der Klimatologie I, Part A, Verlag Borntraeger, Berlin 1930

298. Vernekar, A.D.: Met. Monographs 12, 1 (1971) 299. Berger, A.: Geophysical Surveys 3, 351 (1979) 300. Berger, A.L.: I1 Nuovo Cimento 2C, 63 (1979) 301. Berger, A., et a!.: Long-term variations ofmonthly insolation as related to climatic changes, In

ternational Wegener Symp., Berlin (1980) 302. JOC: The physical basis of c1imate and c1imate modelling, GARP Pub!. Sero No. 16, 5 (1975) 303. ICSU-WMO: Preliminary Plan for the WCRP, ICSU-WMO, WCP-2, Geneva, 1981


304. Cess, R.D.: J. Atmosph. Sei. 33, 1831 (1976) 305. Manabe, S., Wetherald, R.T.: J. Atmosph. Sei. 24, 3, 241 (1967) 306. Ohring, G., Clapp, P.: ibid. 37, 2, 447 (1980) 307. a Shulka, J., Sud, Y.: Effeet of c1oud-radiation feedback on the c1imate of a general eireulation

model, Goddard Lab. Atm. Seiences, NASA/GSFC, Greenbelt, J. Atm. Sei. 38, 2337 (1981) 307. b Stephens, G.L., Webster, P.J.: J. Atmosph. Sei. 38, 235 (1981) 307. e Wetherald, R.T., Manabe, S.: J. Appl. Meteorol. 37, 1485 (1980) 308. Paltridge, G.W.: J. Atm. Sei. 31, 1571 (1971) 309. Wetherald, R.T., Manabe, S.: J. Geophys. Res. 86 C, 1194 (1981) 310. a Manabe, S., Stouffer, R.J.; ibid. 85 C, 5529 (1980) 310. b Gates, W.L. et al.: J. Geophys. Res. 86, C 76385 (1981) 311. a AngelI, J.K., Korshover, J.: Monthly Weather Rev. 106, 755 (1978) 311. cHanen, J., et al., Seienee 213,957 (1981) 312. a Ramanathan, V., Coakley, J.A., Jr.: Rev. Geophys. and Spaee Phys. 16,465 (1978) 312. b Ramanathan, V.: Seienee 190, 50 (1975) 312. e Laeis, A. et al.: Geophys. Res. Let. 8, 1035 (1981) 313. Wang, W.C., et al.: Seienee 194, 685 (1976) 314. Leighton, P.A.: Photoehemistry of Air Pollution, Aeademie Press, New York 1961 315. Leighton, P.A.: Chemieal Reaetions in the Lower and Upper Atmosphere, Interseience Publ.,

New York 1961 316. Hall, T.C., Jr., Blacet, F.E.: J. Chem. Phys. 20, 1745 (1952) 317. Zdunkowski, W.G., Welch, R.M., Paegle, J.: J. Atmosph. Sei. 33, 2399 (1976) 318. a Gerber, H., Hindman, E. eds.: Light absorption by aerosols. Speetrum Press, Hampton, V.A.

1982 318. b Sellers, W.D.: J. App. Met. 12,241 (1973) 319. Chylek, J.P., Coakley, J.A.: Seienee 183,75 (1974) 320. Coakley, J.P., Chylek, P.: J. Atmos. Sei. 32, 409 (1975) 321. Hansen, J.E., et al.: Climate effeets of atmospherie aerosols, Conf. Aerosols: Urban and Rural

Charaeteristies Souree and Transport Studies, New York Acad. Sci. 1979 322. Ohring, G.: PAGEOPH 117, 851 (1979) 323. Temkin, R.L., Snell, F.M.: J. Atmos. Sei 33, 1671 (1976) 324. Charlock, T.P., Sellers, W.D.: J. Atmos. Sci. 37, 1327 (1980) 325. a Toon, O.B., Pollack, J.B.: J. Appl. Meteorol. 15,225 (1976) 325. b Toon, O.B., Pollack, J.B.: Amer. Sci. 68, 268 (1980) 326. Reck, R.A.: Influence of airborne partic1es on the Earth's radiation balance, GARP Pub!. Series

No. 22, p. 947, 1979 327. Reck, R.: Science 186, 1034 (1974), Atmos. Environm. 8, 823 (1974) 328. Wang, W.-C., Domoto, G.A.: J. Appl. Met. 13, 521 (1974) 329. Rasool, S.I., Schneider, S.H.: Science 173, 138 (1971) 330. Yamamoto, G., Tanaka, M.: J. Atmos. Sci. 29, 1405 (1972) 331. Levin, Z., Joseph, J.H., Mekler, Y.: J. Atmos. Sci. 37, 882 (1980) 332. Carlson, T.N., Benjamin, S.G.: ibid. 37, 193 (1980) 333. Shettle, E.P., Fenn, R.W.: Models of the atmospheric aerosols and their optical properties.

AGARD Conf. Proe. No. 183, Optical propagation in the atmosphere, p. 45, 1976 334. Pollack, J.B., et al.: Nature 263, 551 (1976) 335. Pollack, J.B., et al.: J. Geophys. Res. 81, 1071 (1976) 336. Hansen, J.E., Wang, W.C., Lacis, A.A.: Science 299, 1065 (1978) 337. Harshvardhan: J. Atmos. Sci.36, 1274 (1979) 338. Twomey, S.: ibid. 34, 1149 (1977) 339. a Joseph, J.H.: The effect of desert aerosol on a model of the general circulation, Proc. Symp.

Radiation in the Atmosphere [H.-J. Bolle (ed.)], Scienee Press, Princeton 487, 1977 339. b Navato, A.R. et al.: Monthly Wea. Rev. 109, 244 (1981) 339. c Lorius, c.: private communication 340. Kunkel, B., et al.: Air Quality Measurements from Space Platforms, ESA-Report CR-577, 1975 341. Reiter, R.: Meteorolog. Rundschau 28,37 (1975); Wiss. Mitt. Nr. 9, Garm.-Partenk., 1974 342. Herbert, G.A. (Ed.): Geophysical Monitoring for Climate Change No. 8., Summary Rep. 1979,

U.S.Dep. of Commerce, 1980

302 H.-J. Bolle

343. a Cadle, R.D., et al.: Recent studies of the stratospheric aerosollayer, in: Proc. Int. Conf. on Structure, Composition and General Circulation of the Upper and Lower Atmospheres and Possible Anthropogenic Perturbations, IAMAP, Toronto 1974

343. b White, D.E., Waring, G.A.: U.S. Geo!. Survey Prof. Paper 440-K, 1963 344. Hobbs, P.V., et a!.: Science 211,816 (1981) 345. Inn, E.C., et al.: ibid. 211, 821 (1981) 346. Gandrud, B.W., Lazrus, A.L.: ibid. 211, 826 (1981) 347. Murcray, D.G., et al.: ibid. 211, 823 (1981) 348. Vossler, T., et a!.: ibid. 211, 827 (1981) 349. Zoller, W.H., et a!.: Trace metals in the Antarctic atmosphere, in: WMO: Observation and mea

surement of atmospheric pollution, Spec. Envir. Rep. No.3, WMO-No.368, 380-386, Geneva 1974

350. Farlow, N.H., et al.: Science 211, 832 (1981) 351. Petterson, E.M.: ibid. 211, 836 (1981) 352. Chuan, R.L., Woods, D.C., McCormick, M.P.: ibid. 211, 830 (1981) 353. McCormick, M.P., et a!.: Bull. Amer. Meteoro!. Soc. 60, 1038 (1979) 354. McCormick, M.P. et a!.: Science 214, 328-331 355. a McCormick, M.P.: SAM II measurements of the polar stratospheric aerosol, Vo!. 1., NASA

Ref. Pub!. 1081 (1981) 355. b McCormick, M.P. et al.: SAGE measurements ofthe stratospheric aerosol dispersion and load-

ing from the Soufriere volcano. NASA Technical Papers 1922 (1981) 356. Budyko, M.l.: Tellus 2, 611 (1969) 357. Lian, M.S., Cess, R.D.: J. Atmosph. Sci. 34, 1058 (1977) 358. Ghil, M., Bhattacharya: An energy-balance model of glaciation cyc1es, in: Rep. JOC Study Conf.

Climate Models: Performance, Intercomparison and Sensitivity Studies, WMO-ICSU, GARP Pub!. Series No. 22, 1979; World Meteoro!. Organization: Observation and measurement of atmospheric pollution. Spec. Environ. Rep. No. 3, WMO-No. 368, Geneva 1974

359. Fletcher, J.O. (Ed.): Proc. Symp. Arctic Heat Budget and Atmospheric Circulation, Memorandum RM-5233-NSF, Rand Corp. 1966

360. Untersteiner, N.: J. Geophys. Res. 69, 4755 (1964) 361. Badgley, F.l.: Heat balance at the surface of the Arctic Ocean, Proc. Symp. Arctic Heat Budget

and Atmospheric Circulation [J.O. Fletcher (ed.)], Rand Corp., p. 215, 1966 362. Herman, G.F., Johnson, W.T.: Monthly Weather Rev. 108, 1974 (1980) 363. Weller, G.: ibid. 108, 2006 (1980) 264. a Viebrock, H.: J. Geophys. Res. 67, 4293 (1962) 364. b Zillman, J.W.: Antarctic Res. Ser., Amer. Geophys. Union 19, 11 (1972) 365. Bunker, A.F.: Monthly Weather Rev. 104, 1122 (1976) 366. Winston, J., et a!.: Earth-atmosphere radiation budget analyses derived from NOAA satellite

data, June 1974--February 1978, U.S.Dept. Commerce, Meteoro!. Satellite Laboratory, p. 343, 1979

367. Maykut, G.A.: J. Geophys. Res. 83, 3646 (1978) 368. Gavrilova, M.K.: Radiation Climate ofthe Arctic, Leningrad, Gidormeteorologicheskeo Izd., 178

pp., 1963 (English trans. by Israel Prog. for Sci. Trans!. 1966) 369. Dalrymple, P., Lettau, H., Wollaston, S.: South Pole micrometeorological program: data analysis.

Antarctic Res., Geophys, Monogr. No.9, Amer. Geophys. Union, p. 13, 1966 370. Schwerdtfeger, W.: Antarct. J. (U.S.) 3, 193 (1968) 371. Liljequist, G .H.: Energy exchange of an antarctic snowfield. Norwegian-British-Seedish Antarctic

Expedition, 1949-52, Sci. Res., Part I, Vo!. II, Oslo, 1956 (Available from Norsk Polarinstitutt, Postboks 157, N-1330 Oslo Lufthavn, Norway)

372. Rusin, N.P.: Meteorological and Radiational Regime of Antarctica (Russian) Trans!. by the Israel Program for Sei. Translations, Washington, DC, p. 335, 1961

373. Weller, G.: J. Gcophys. Res. 73, 1209 (1968) 374. Lemke, P., Trinkl, E.W., Hasselmann, K.: J. Phys. Oceanography 10, 2100 (1980) 375. Ottermann, J.: Science 86 531 (1974) 376. Chamey, J.: Dynamics of desert and drought in the Sahei, in: The Physical Basis of Climate and

Climate Modelling, JOC-GARP Pub!. Ser. 16, 171 (1976) 377. Seginer, I.: Agric. Meteorology 6,5 (1969)


378. Idso, S.B., et al.: J. Appl. Meteorology J 4, 109 (1975) 379. Kukla, G., Robinson, D.: Monthly Weather Rev. 108, 56 (1980) 380. Global Monitoring of the Environment for Selected Atmospheric Constituents, joint Publication

of the World Meteorolog. Organisation, The Environm. Protection Agency and the U.S.Deptm. of CommercejNOAA in co-operation with the United Nations Environment Programme, prepared by Environmental Data and Information Service, National Climatic Center, Asheville, USA

381. Stern, A.C. (ed.): Air Pollution, Vol. III, Academic Press, New York, San Francisco, London 1976 382. Lundgren, D.A., et al. (eds.): Aerosol Measurement, Univers. Press Florida, Gainesville 1979 383. WMO: Spec. Environment Rep. 3, WMO Publication, p. 368, 1974 384. WMO: Spec. Environment Rep. 10, WMO Publication, p. 460, 1976 385. a Komhyr, W.D., Harris, T.B.: Measurement of atmospheric CO2 at the USA GMCC baseline

stations, Proc. WMO Air Pollution Measurement Techniques Conf., 11-15 Oct. 1976, Gothenburg, Sweden, WMO Publ. 460, Spec. Environ. Rep. 10, 9 (1976)

385. b WMO: Environmental Pollution Monitoring Programme, Summary Report on the Status of the WMO Background Air Pollution Monitoring Networks as at April 1981, Geneva, 1981

386. WMO: Techn. Conf. Regional and Global Observations of Atmospheric Pollution Relative to Climate, Spec. Environment Rep. 14, WMO Publication, p. 549, 1979

387. Komhyr, W.D., Thompson, T.M., Dutton, E.G.: Chlorofluorocarbon-II, -12. and nitrous oxide measurements at the U.S. GMCC baseline stations (16. Sept. 1973 to 31. Dec. 1979), NOAA Techn. Rep., 1981

388. DeLuisi, J.J.: J. Geophys. Res. 84 C, 1766 (1979) 389. DeLuisi, J.J.: Appl. Optics 183190 (1979) 390. DeLuisi, J.J., Mateer, c.L., Heath, D.F.: Geophys. Res. 84 C, 3728 (1979) 391. Pollack, J.B., McCormick, M.P. (Eds.): Special Issue on Aircraft and Spacecraft Measurements

of Stratospheric Aerosols and their implications, Geophys. Res. Lett. 8, 2 (1981) 392. a Fabian, P., et al.: J. Geophys. Res. 84, 3149 (1979) 392. b Fabian, P.: Atmospheric sampling, in: Adv. in Space Res. 1, No.ll, p. 17, Pergamon Press, Ox-

ford 1981 393. a Goldan, P.D., et al.: J. Geophys. Res. 85C, 413 (1980) 393. b Farmer, c.B., et al.: ibid. 85 C, 1621 (1980) 393. c Niple, E., et al.: ibid. 7,489 (1980) 394. Gille, J.C., RusselI, J.M., Bailey, P.L.: ibid. p. 267 395. Russell III, et al.: Adv. Space Res. 1, First FGGE Results from Satellites, Pergamon Press, p. 27,

1981 396. Fischer, H., Gille, J., RusselI, J.: Adv. Space Res. 1, First FGGE Results from Satellites, Perga-

rnon Press, p. 279, 1981 397. Taylor, F.W., et al.: ibid., p. 261 398. Baker, D., Steed, A., Stair, A.T. Jr.: Appl. Optics 20,1734 (1981) 399. Twomey, S.: Induction to the Mathematics ofInversion in Remote Sensing and Indirect Measure-

ments, Development in Geomathematics 3, Elsevier, Amsterdam 1977 400. Hinkley, E.D. (Ed.): Laser Monitoring of the Atmosphere, Springer, New York 1976 401. Uthe, E.E.: Appl. Optics 20, 1503 (1981) 402. Mörl, P., et al.: Contributions to Atmospheric Physics 4, 403 (1981)

Subject Index References of Part Aare marked A, References of Part Bare marked B.

Abietic acid B: 124 Absorption by aerosols B: 154 - band of water B: 183 - bands B: 161 - coefficient B: 140, 168 - -, water vapor B: 163 - number B: 152, 154 - in particles B: 151 - processes B: 153 Acer saccharum, nutrient quantities B: 7 Acid rain A: 42 Actinide decay series B: 51 Actinides B: 48 Actinium decay series B: 55 Acyclic isoprenoid hydrocarbons B: 113 Adiabatic lapse rate B: 238 - temperatures profile B: 227 Adenosine-3',5'-monophosphate A: 148 Adenosine triphosphate A: 147 Adsorption see sorption -, oxide surfaces A: 37 Advective fluxes B: 207 Aedes aegypti B: 5 Aerobic respiration A: 91; B: 105 Aerological stations B: 283 Aeroplysinin-I A: 242 Aerothionin A: 242 Aerosol albedo feedback B: 278 -, background components B: 265 -, climate effect B: 261, 262 - Climate Effects pro gram B: 285 -, infrared B: 262 - at Mauna Loa B: 266 -, optical properties B: 261, 262 - scattering B: 150 - size distributions B: 152; B: 153 - at South Pole Station B: 266 -, stratospheric B: 263 -, - background B: 263 -, tropospheric B: 261 - turbidity values B: 175 Aerosols B: 133, 177, 238

-, continental B: 261 -, maritime B: 261 -, organic nitrogen compounds B: 77 -, volcanic B: 258, 265 Age distribution B: 13 - pyramid, microtus agrestis B: 13 Agrostos tenuis B: II Air constituents, vertical distribution B: 283 -, ecology B: 9 Airglow B: 169 Albedo B: 142, 154, 209, 255, 261 - annual zone means B: 216 - of clouds B: 215 -, dependence B: 199 - of a glacier B: 199 -, planetary B: 209 - of sea surface B: 192 - of soil B: 279 - on solar zenith angle B: 199 -, solid earth B: 194 - spectral of natural surfaces B: 198 Aigae A: 161 Aigal metabolites A: 240 n-Alkanes, in geochemistry B: 112 Allelochemics B: 7 Allelopathic effects B: 8 Allogenic B: 28 Allopatry B: 18 Alpha-diversity B: 21 Alpine biomes B: 35 Aluminium silicates A: 43 Ammonia B: 63, 161 - cycle B: 73 Angasiol A: 246 Angular distribution, reflected radiation

B: 194 Anisotropy factor B: 147 Antarctica B: 275 -, radiation budgets B: 221 Anthropogenic climaxes B: 30 - materials in ocean A: 65 "Anticyclonic" curvature B: 242

306

Aplysiatoxin A: 231, 246 Aplysin A: 245 Approximation by spherical harmonics B:202 Approximative similarity solution B: 203 Aquatic ecosystem B: 23 - system, partitioning ofthe global inventories of

nitrogen B: 71 Aqueous carbonate system A: 32 - reactions of sulfur A: 126 Archaebacteria B: 114 Arsenic, biological cyde A: 219 -, methylation A: 217 Artemisia B: 35 Astronomical Unit B: 171 Asymmetry factor B: 151 Atlantic circulation system B: 246 Atmosphere A: 106, 107; B: 134, 173 -, composition A: 2 -, energy budget B: 249 -, flux of sulfur A: 115 -, mercury pool A: 180 -, natural radio-nudides B: 55 -, origin A: 4 -, ocean B: 135 -, optical depth B: 176 -, scale height B: 176 -, transmittance B: 176 Atmosphere-earth system B: 261 Atmosphere-ocean system B: 240 Atmospheric band system B: 169 - carbon dioxide content, environmental

responses B: 102 - constituents B: 161, 252 - emission B: 183 - -, non-thermal B: 170 - emissions, metals A: 46 - gases B: 160 - oxygen A: 89 - reactions of sulfur A: 125 - spectrum B: 181 - transmittance B: 174 - zenith emission spectrum B: 184 Aurora B: 169, 239 Australian mallee B: 34 Autecology B: 1 Autogenic B: 28 Autotrophs B: 4

Bacteriochlorophyll B 94 Balloons B: 285 Band absorption B: 157; B: 158 - transmissivities B: 206 Banded iron formation A: 13 Baseline Pollution Stations B: 283 Beer's law B: 140 Behavioral Relationships, ecology B: 8 Benz(a)anthracene B: 125 Benzo(a)pyrene B: 125

Benzopyrene, asbestos B: 125 Bergmann's rule B: 10 Beta-diversity B: 22 Biochemical fossils B: 113

Subject Index

- redox cyde A: 40 Bioconcentration, of metals A: 210 Biogeochemical nitrogen cyde B: 61 - - -, microorganisms B: 62 Biological activity in seawater A: 54 - CO2 cycle B: 104 - cycle B: 84 - -, carbon B: 85 - markers B: 117 - methylation A: 169ff. - sulfur compounds A: 128 - transformations of sulfur A: 128 Biomass B: 4 Biomes B: 3, 33 Biomethylation A: 170 - of heavy metals A: 214 -, lead A: 195 - of mercury A: 183 Biosynthesis, halogenated natural products

A: 248 Biotic ecology B: 2 - resources, exploitation B: 41 Black body emission B: 166 - - excitance B: 167 Blue color of the sky B: 147 Boreal forest, primary production B: 25 Bottom boundary layer B: 137 Bowen ration B: 222 Brominated phenols A: 239 5-Bromo-3,4-dihydroxybenzaldehyde A: 239 5-Bromo-N,N-dimethyl-tryptamine A: 244 6-Bromohypaphorine A: 244 "Bulk parameter" B: 145 Buyancy B: 208

Cadmium A: 212 -, biomethylation A: 215 Calcite precipitation B: 97 Calcium phosphates, minerals A: 152 Californian chaparall B: 34 - redwoods B: 34 Calluna B: 28 Camberlain bands B: 169 Canopy B: 20 Carbon balance B: 86 - -, ocean B: 89 - burial rate A: 100 - budget, global B: 84 - consumption capacity A: 99 - cycle B: 83; B: 84 - -, global B: 86 - -, reservoirs B: 85 - -, residence time B: 85

Subject Index

Carbon dioxide A: 7; B: 4, 83, 160, 162,239 - - assimilation B: 95 - -, atmospheric reservoir A: 7 - - bands B: 162 - -, biological cycle B: 203 - -, climate impact B: 257 - -, infrared absorption A: 3 - - problem B: 98 - -, regulation B: 102 - -, in seawater A: 55 - -, sinks B: 99 - -, sources B: 99 - disulfide in seawater A: 111 - isotope mass balance A: 9 - isotopes, stable A: 101 Carbon-13, mass balance A: 10 Carbon monoxide A: 93; B: 160 - - in seawater A: 55 Carbon-14 in oceanography A: 54 Carbon preference index (C.P.I.) B: 112 - tetrabromide in marine organisms A: 231 - tetrachloride in marine organisms A: 231 Carbonate A: 30 - deposition B: 92 - equilibria A: 26 Carnivores B: 23 CH4 see Methane Chamaephytes B: 11 Chaparral shrubs B: 8 Chappius continuum B: 164 Chelate A: 34 Chemical Relationships, ecology B: 8 - weathering A: 43 Chemotaxonomy A: 229 Chernozem B: 34 Chlorinity A: 52 Chlorofluoromethane A: 3 -, ozone destruction A: 3 Chloroform in marine organisms A: 231 Chlorophyll B: 93, 95, 113 -, absorption B: 94, 194 -, photoactions B: 94 Chondriol A: 239 CIO B: 162 Climate B: 240 - change B: 252, 270 -, ecology B: 5 - parameters, monitoring B: 279 - Research B: 252 - sensitive parameter B: 270 - variations B: 133, 253

Climax concept, ecology B: 28 - pattern B: 29 Cloud albedo B: 255 - - feedback B: 262 - cover B: 133 - coverage B: 217

Cloudiness B: 215 Clouding of lakes B: 98 Cloud-radiation feedback B: 255; B: 256 Cloud, white appearance B: 183 Clouds B: 238, 254, 263 -, absorption B: 186 -, albedo B: 186; B: 188 -, climatic impacts B: 254 -, polluted B: 189 -, radiation properties B: 183fT. -, transmittance B: 186 -, water content B: 185 CO see carbon monoxide CO2 see carbon dioxide - flask sampling B: 284 - vibration-rotation band B: 159 Coal A: 78 Coastal zones A: 56 Cobalamin A: 172 Coenocline B: 22 Cohorts B: 14 Competition, ecology B: 12, 17 Competitive exclusion principle B: 18 Complexation chemistry A: 33 fT. Complexes, clay-metal-organic A: 79 - of metals A: 208 Complex-gradient B: 22 Communities, ecology B: 19 -, level B: 2 -, terrestrial B: 20 Concinndiol A: 238 Condensation nuclei B: 263

307

Continental earth's crust, natural radionuclides B: 55

- red beds A: 14

Continuum absorption of water vapor B: 163 Convection layer, atmosphere B: 137 Convergence, ecology B: 11 Cooling of earth surface B: 210 - rate of atmosphere B: 249 Copepods B: 114 Coral reefs B: 39 Coriolis force B: 134 Corona B: 173 Cosmogenic radionuclides B: 50, 55 Critical climate parameters B: 256 Cromosphere B: 173 Crude oil B: 112 - oils, alkylhomologues of PAH B: 125 Crustal abundance, some key elements A: 204 Cryptophytes B: 11 Cryptozoic animals B: 20 Cultivated land, net primary production B: 87 - -, primary production B: 25 Cyanobacteria B: 93, 96 Cycle of arsenic A: 219 -, biochemical B: 26

308

Cycle, biogeochemical B: 27 - of lead A: 185 - of manganese A: 213 - of mercury A: 177 - of oxygen A: 87 ff. - of phosphorus A: 147 ff. - of sulfur A: 105ff. Cycles of metals A: 169 ff. - - other than Hg, Pb and Sn A: 202 - of tin A: 197 Cyclic photophosphorylation B: 93 Cycling, ecology B: 27 Cyclomopol A: 231 Cyclonic curvature B: 242

Damping depth for heat wave B: 224 Danaus plexippus B: 8 Day B: 214 DDT A: 66 Deep sea sediments A: 61 Deer mouse, densities of the cocoon B: 18 Demethylation processes A: 177 Denitrification B: 63, 66, 103, 105 Deposits, of phosphate A: 153 Desert, net primary production B: 87 -, primary production B: 25 Deserts B: 35, 262 Desoxy Hopanoids B: 121 Detergent phosphate A: 162 Detrital sulfides A: 13 Detritus A: 80 Dialkyl tin compounds, stabilizer A: 202 5,6-Dibromo-N,N-dimethyl-tryptamine A: 244 Dibromophakellin A: 243 4,5- Dibromopyrrole-2-carboxylic acid A: 243 Didinium nasutum B: 17 Diffuse solar radiation B: 209 Dimethyl mercury A: 177 - sulfide A: 93, 118 Disclimax B: 29 Discret space theory B: 203 Dispersal, ecology B: 13 Dispersion, ecology B: 16 Dissipation rate for kinetic energy B: 235 Dissolved organic carbon (DOC) B: 90 Dissociation B: 143 - wavelength treshold B: 165 Dissociative absorptions B: 164 Diterpenoids B: 124 Diversity-stability-hypothesis B: 31 Dodecaton approach to radiative transfer

B: 203 Dominance, ecology B: 21 Doppler line broadening B: 155 Doubling adding method B: 203 Drag coefficient B: 227 Drop size distribution B: 184

Dust A: 119 Dysidin A: 241

Earth-atmosphere system B: 131

Subject Index

- -, annual zonally energy fluxes B: 250 - -, energy budget B: 247 Earth orbit B: 213 Earth's crust, composition A: 70 E-I-chlorotridec-I-ene-6,8-diol, algae A: 231 Ecological formations B: 3 - isolation B: 18 - systems B: 2 Ecology B: I -, feedbacks B: 30 Ecophysiology B: I Ecosystem types B: 25 Ecosystems B: 23 -, development B: 29 -, freshwater B: 36 -, level B: 3 -, marine B: 37, 38 -, terrestrial B: 32 Ecotope B: 23 Ecotypes B: II Eddies B: 235 Eddington approximation B: 203 Eddy correlation method B: 227 - diffusivity B: 227 Eddys, transient B: 242 Ekman layer B: 134 - pumping B: 244 Electrical field strength B: 145 Electromagnetic spectrum, sun B: 143 Electron excitation B: 143 - gas heating rate B: 239 Electronic spectra B: 164 Element concentrations, St. Helens eruption

B: 269 Elements, crustal abundance A: 205 Eltonian pyramids B: 24, 25 Emitted flux annual zonal means B: 218 Emission, infrared B: 132 Emissivity method B: 207 -, natural surfaces B: 197, 199 - of water B: 191 Endothermic animals B: 4 Energy B: 235, 236 - balance, different surfaces B: 276, 277 - budget B: 209 - - of earth-atmosphere system B: 207 - conversion, photosynthesis B: 93 - conversions B: 132, 208 - deposition in the atmosphere B: 236 -, ecology B: 4 - flow, ecology B: 27 - fluxes B: 207 - -, atmosphere B: 212, 231

Subject Index

- - at the earth surface B: 220 - sources and sinks, vertical distribution B: 237 - transport B: 131 - transports, extreme values B: 234 Enrichment factors, metals A: 206 Entropy production B: 132 Environment, abiotic, ecology B: 4 -, biotic B: 7 Environmental adaptation B: 10 - concentrations of lead A: 188 - disruption B: 42 - gradients of communities B: 22 - problems, pollution B: 40 - transport A: 199 - - of lead A: 188 - -, sulfur A: 115 Epilimnion B: 36 Epiphytes B: 20 Equation of radiative transfer B: 201 ff. Equitability, ecology B: 21 Erigone arctica B: 14 Estuaries B: 38 Eukaryotes A: 14 Eukaryotic algae B: 96 - organisms B: 115 Euphotic zone B: 38, 90, 137 Euryhaline B: 9 Eurythermal B: 9 Eutrophication A: 162 Evaporation B: 228, 230 Evolutionary changes, systems B: 31 Exchange processes, natural radionuclide

tracers B: 59 Excitance B: 166 Exosphere B: 136 Exponential Kemal Approximation B: 204 Extinction coefficient B: 140

Feedback process B: 133 Feedbacks, negative, ecology B: 30 Feeding Relationships, ecology B: 8 Feldspars A: 73 Fermentation B: 105 Fluoranthene, manganese nodules B: 125 -, mercury ores B: 125 Fluoroapatite A: 152 Flux across latitude B: 232 Food chains B: 24 - -, saprothroph B: 24 - -, heterotroph B: 24 -, use by animals B: 26 - web B: 24 Forest brown earth B: 34 Forests, net primary production B: 87 -, temperate B: 33 -, tropical B: 33 Fossil fuel buming B: 100

- - combustion A: 118 - -, release of natural radionuclides B: 57 Fraunhofer lines B: 171 Fresh waters A: 41 Functional response, ecology B: 17 - unit, ecology B: 23 "Fulvic Acids" B: 88 Fundamental niche B: 12

Gasterosteus aculeatus B: 11 Gauss-Seidel iteration B: 203 Gazing animals B: 23 Gelbstoffe A: 56 Genotypic B: 10 Geochemical classification of the elements

A: 70 - processes B: 47 - uniformitarianism A: 11 Geochemistry, organic B: 111 Geological classification, soils A: 72 - cycle B: 84 - -, carbon B: 85 Geopotential height B: 231 Glaciation B: 270 - periods B: 252 Glaciations B: 253 Global interference factors A: 205 - nitrogen cycle B: 78 - radiation B: 187, 209 - Teletype System B: 283 - warming B: 100 Grasslands B: 35 Gravity field A: 4

309

Greenhouse effect A: 6, 109; B: 100, 255, 256, 260,262

Green sulfur bacteria B: 96 Growth efficiency B: 26 Guano A: 153 Gypsum A: 75, 107

Habitat, ecology B: 11 Hadley cell B: 242 Halomethanes in Asparagopsis species A: 232 Hazy atmosphere, Scattering B: 152 HBr B: 161 HCI B: 161 H2CO B: 162 HDO B: 162 Heat budget atmosphere B: 249 - -, ice and snow surfaces B: 270 - -, polar areas B: 270, 271 - capacity, specific B: 223 - exchange between oceans and atmosphere

B: 246 - flux B: 209 - - at earth surface B: 224 - - into the ground B: 223 - - from interior of the earth B: 222

310

Heat fluxes, aretie regions B: 274 - -, latitudinal distribution B: 249 Heating of earth surfaee B: 210 Heat transports B: 263 Heavy metals, biomethylation of A: 214 - -, toxieity A: 204 Hemieryptophytes B: 11 Herbs B: 20 Herzberg bands B: 169 - eontinuum B: 164 Heterotrophs B: 4 Hexabromobipyrrole A: 230 Hexose-l,6-diphosphate B: 95 HF B: 161 HI B: 161 HN03 B: 161 HzO B: 160, 162 Honeybee workers, life table B: 14 Hopanepolyols, diagenesis B: 122 C35-Hopanepolyols B: 122 Hopanoid skeleton B: 118 Human population growth B: 42 Humie acids A: 55; B: 88 "Humins" B: 88 Hydrocarbons A: 93 Hydrogen bonding A: 19 - loss to spaee A: 4 - sulfide A: 93, 96 - - in the atmosphere A: 107 Hydrologic cycle A: 41, 47 Hydrosphere A: 17 ff., 106 -, flux of sulfur A: 121 -, sulfur A: 110 Hydroxyapatite A: 152 Hyperparasites B: 24 Hypogastrura viatica B: 14 Hypolimnion B: 36

Ice A: 19 Ice absorption bands B: 197 - age B: 252 lee-albedo feedback B: 270 Ice clouds B: 187 - and snow cover B: 278 - - - of the earth B: 273 -, vapor pressure A: 23 Individualistic hypothesis B: 22 Individuals, eeology B: 4 Infrared emission of the atmosphere B: 220 - excitance B: 217 - flux density B: 219 Ingenous rocks A: 71 Inner boundary layer B: 134 Inositol phosphates A: 156 Interfaciallayer B: 137 Interference indices A: 206

Subjeet Index

International Association of Meteorology and Atmosphere Physies B: 140

- Satelite Cloud Climatology Project B: 256 Interspecific competition, ecology B: 18 Interstitial fauna B: 38 Intertidal zone B: 38 "Inter-Tropieal Convergence Zone" B: 241 Intraspecific eompetition, eeology B: 18 Invariance principle B: 203 Invariat imbedding method B: 203 Inventories of nitrogen B: 68 Ionisation B: 143 Irieol A: 238 Iran reduetion B: 105 Irradiance B: 138 Irradiances under cloudy sky B: 189 Isocaespitol A: 235 Isolaureatin A: 238 Isomaneonene A: 239 Isoprenoids B: 114

Jet streams B: 242, 243

Kaolinite A: 43 Kaplan-Meinel bands B: 169 Kerogen A: 10, 95, 101 Key factor, eeology B: 16 Kirchhoffs law B: 166

Lake and stream, net primary production B: 87 Lakes, stratification B: 20 Laminar sublayer B: 134 Lanasol A: 239 Lapse rate B: 134 - -, adiabatic B: 136 Latent heat flux B: 228, 230 - - fluxes B: 210 Laureatin A: 238 Laurefuein A: 238 Lautrinterol A: 235 Lead, anthropogenie losses A: 189 - in an aquatic Eco-system A: 187 -, biogeochemieal cycles A: 185 -, crustal abundanee A: 186 -, environmental coneentration A: 186 -, global cycle A: 194 -, methylation A: 195 - m oeeans A: 191 - particles from gasoline A: 190 - in soils and sediments A: 192 Lecithin A: 148 Leguminosae B: 64 Lentic communities B: 36 Lewis acid A: 34 Lianes B: 20 Lidar B: 285 Life cycle B: 15 - history B: 15

Subject Index

Light, ecology B: 9 Lightning A: 94 Limiting factors, ecology B: 8, 16 Limnetic zone B: 20 Line absorption B: 155 Line absorption coefficient B: 155, 156 Liquid water mass B: 183 Lithosphere A: 69, 106 -, flux of sulfur A: 122 -, sulfur content A: 113 -, sulfur oxidation A: 126 Littoral zone B: 20 Local thermodynamic equilibrium B: 166, 183 Longwave infrared B: 194 - radiation B: 180,215 Lotic communities B: 36, 37 Lorentz line shape B: 155 Lyman-Birge-Hopfield bands B: 164 Lyman-a-emission of hydrogen B: 164

Macronutrients B: 6 Magnetosphere B: 136 Mammals, oral toxicity of metals A: 211 Maneonene A: 239 Manganese A: 212 - cycles A: 213 - reduction B: 105 Marine algae, organohalogen compounds

A: 230 - animals, tin A: 198 - ecosystems, primary production B: 91 - -, standing crop B: 91 - environment, natural radionuclides B: 55 - invertebrates, organohalogen compounds

A: 245 - natural products A: 229, 249 - sediment, natural radionuclides B: 55 Mauna Loa B: 100 Maxwell's theory B: 144 Mean meridional circulation B: 235 Mercury, anthropogenie losses A: 182 -, biogeochemical cycles A: 177 - cycle, aquatic A: 179 - methylation, photochemical A: 176 - in natural waters A: 181 Meridional circulation B: 232 - fluxes B: 211 Mesopause B: 136 Mesophere B: 136,239 Metal cycles A: 169 ff. - Ions A: 35 Metalloids, biomethylation A: 217 Metals, atmospheric emissions A: 46 -, biogeochemical cycle A: 171 -, c1assification A: 207 -, oral toxicities to small mammals A: 211 -, stability of complexes A: 208

Metamorphie rocks A: 71 Meteorological equator B: 241, 246 Methane A: 92; B: 160 - formation B: 105 Methanogenic bacteria A: 92 Method of discrete ordinates B: 202 - of successive orders of scattering B: 203 Methy1chrysenes B: 125 Methyl cobalamin A: 172, 173 - iodide A: 250 - - in marine organisms A: 231 - mercury in fish A: 181 Methylation see biomethylation -, biological A: 169ff. - of metals A: 215 Methylmercaptane A: 118 Methylmercurythiomethyl A: 184 Micas A: 73 Microbial fermentation A: 92 - transformation, triterpenoids B: 120 Microclimate B: 5 Micronutrients B: 6 Microorganism involved in sulfur reactions

A:134 Microorganisms, redox reaction A: 39 Microtus agrestis B: 13 Microwave radiation B: 141, 143 Mie theory B: 145, 150 Minera1ization B: 63 -, nitrogen B: 65 Minerals, in soils A: 73 Mining, lead emission A: 191

311

Minor trace constituent concentrations B: 260 Mixing processes B: 58 Molecular rotation B: 143 - vibration-rotation B: 143 Monarch butterfly B: 8 Monazite B: 54 Monitoring, upper atmosphere B: 285 Monoclimax B: 29 Monosoons B: 243 Monte Carlo method B: 202 Mortality, ecology B: 13 Mosquito B: 5 Mount St. Helen B: 267, 268 Multidimensional niche B: II Multiple scattering B: 183 Multipoles B: 150 Muscovite A: 73

Namibian Shelf sediments B: 126 Natality, ecology B: 13 Natural radioactive isotopes B: 53 - radionuclides B: 53 Negentropy B: 92 Nekton B: 20 Neodiprion, densities of the cocoon B: 18

312

Neritic Ecosystems B: 39 Net primary production, cultivated land B: 87 - - -, desert B: 87 - - -, forests B: 87 - - -, lake and stream B: 87 - - -, savana B: 87 - - -, swamp and marsh B: 87 - - -, temperate grassland B: 87 - - -, tundra B: 87 - radiation flux B: 219 - - - densities B: 220 Network for Monitoring Background Air

Pollution B: 279, 282 NH 3 see ammonia Niche, ecology B: II Nitrate assimilation A: 94 - ammonification B: 105 Nitrification B: 63, 65

Nitrogen A: 8 - absorption B: 164

in the atmosphere, Partitioning of the various forms B: 70 budgets B: 78 compounds, chemical transformations B: 62 cycles B: 61 dioxide B: 63 fixation A: 160; B: 62, 63

- -, organisms involved B: 64 -, global fluxes B: 73 -, in yen tories B: 68 -, -, atmospheric B: 69 -, -, global B: 69 - metabolism, Nitrosomonas B: 66 - -, nutrients B: 10 - oxides, ozone destruction A: 3 - transformation, abiotic B: 67 Nitrosococcus B: 65 Nitrosolobus B: 65 Nitrosomonas B: 65 -, nitrogen metabolism B: 66 Nitrosospira B: 65 Nitrosovibrio B: 65 Nitrous acid B: 63 - oxide B: 63 - -, in seawater A: 55 - -, sink B: 76 NO B: 161 N 20 B: 160 Noble gases A: 5 N2 0 cycle B: 74, 75 Noncyclic photophosphorylation B: 95 Non-essential elements B: 6 Non-thermal emissions B: 169 Nuclear power B: 57 Nucleotides A: 156 Nuclides, primordial B: 50 Numerical response, ecology B: 17

Nutrient cycling B: 27 - cycle A: 38 -, ecology B: 9, 27

O2 see oxygen 0 3 see ozone Oak B: 5

Subject Index

Obtusadiol A: 238 Ocean-atmosphere system B: 208 ff. Ocean, carbon balance B: 89 - currents B: 245 -, fate of the carbon B: 90 Oceanic ecosystems B: 39 - heat transport, annual mean B: 247 - sediments A: 96 Ocean, penetration of shortwave radiation

B: 191 -, phosphates A: 164 Oceanography, chemical A: 51 ff. Oceans A: 56 -, energy transport B: 243 -, heat transport B: 246, 248 -, lead input A: 191 -, mixed layer A: 48 -, storage and release CO2 B: 257 Ochtodene A: 234 OCS B: 162 OH B: 161 Open ocean, primary production B: 25 Oppositol A: 237 Optical depth of the atmosphere B: 176 Oregonene A A: 234 organic carbon, oxidation A: 95 - - of seawater A: 53 - -, transformation in water A: 40 - geochemistry B: 111 - material, in seawater A: 64 - matter in soil A: 76 - -, soil turnover A: 80 - nitrogen compounds B: 77 - - -, aerosols B: 77 - - transfer B: 77 Organism, level B: 2 Organization levels B: 2 - -, ecological systems B: 2 Organo tin, biocide A: 199 - - compounds A: 199 Organohalogen compounds, natural A: 229 Oroidin A: 243 Oxidation of organic matter B: 105 Oxygen B: 160 -, annuallosses A: 91 -, atmospheric A: 101 -, biochemical cycles A: 89 - budget A: 9 - consumption A: 90 - cycle A: 87 ff.

Subject Index

- demand A: 40 - green line, atomic B: 169 -, origin A: 8 -, photochemical origin A: 8 -, photosynthetic A: 12 3-0xy-triterpenoids B: 119 Ozone A: 3, 94; B: 160, 164, 238, 258 -, destruction A: 3 -, formation A: 3 - layer B: 133 - monitoring B: 285

Pacifenol A: 236 Pack ice B: 273 PAH B: 125 -, sediments B 125 Paramecium caudatum B: 17 Parameterization of band structuren N 206 Particulate organic carbon (POC) A: 59; B: 90 Partitioning of Energy at the surface B: 230 Pathway processes, natural radionuc1ide tracers

B: 59 PCB's A: 66 Pedosphere A: 106 -, fiux of sulfur A: 119 -, sulfur compounds A: 109 Perforatone A: 236 Peroxyacyl nitrate B: 63 Perturbation method B: 204 Perylene B: 126 -, recent sediments B: 125 Perylenequinone pigments B: 126 Phanerophytes B: 11 Phenotypic B: 10 Phosphate biocyc1e A: 160 -, calcium minerals A: 152 - deposits A: 153 - minerals A: 151 - in water A: 157 Phosphates A: 148 Phospholipids A: 156 Phosphoric acid A: 147 Phosphorus cyc1e A: 147ff. -, natural abundance A: 149 Phosphorylation A: 158 Photochemical conversion, triterpenoids B: 120 Photo-chemistry B: 141 Photo dissociation process of atmospheric

species B: 165 Photo-ionisation B: 154, 164, 239 Photolysis of water A: 91 Photosphere B: 171 Photosynthesis B: 89, 92, 209, 229 -, nutrients B: 10 - reaction, global A: 9 -, relation with light intensities B: 10 Photosynthetic organisms A: 87; B: 96

Phototropic bacteria B: 93 Physiognomy B: 19 Phytane B: 113 Phytol B: 113 Phytoplankton A: 57, 230 Picene, mercury ores B: 125 Planetaryalbedo B: 214--216, 219 - boundary layer B: 134 Plankton A: 163; B 20 Plant uptake, sulfate sulfur A: 121 Plants, n-alkanes B: 112 -, phosphate needs A: 159 Podzols B: 34 Polar caps B: 243 - energy budget B: 240 - fronts B: 242 Polarization B: 146 Pollutants B: 279 Pollution, anthropogenie B: 264 -, ocean A: 65 - and ecology B: 40 - monitoring B: 174 Polybrominated biphenyl ethers A: 242 Polycyc1ic aromatic Hydrocarbons, in

geochemistry B: 125 Population sizes B: 15 Populations, ecology B: 12 -, growth form B: 15 -, level B: 2 -, relationships B: 13 Potassium, Isotopes B: 48 "potential evaporation" B: 228 Poynting vector B: 145, 147, 237 Prandl layer B: 134 Precambrian A: 12 Predator-prey relationships B: 17 Preintricatol A: 235 Primary consumers B: 24 - producers B: 24 - production, global B: 25 - productivity, ecosystem B: 25 Processes, geochemical B: 47 Producers B: 23 Profundalzone B: 20 Prokaryotic Red. organisms B: 115 Proton reduction B: 105 Purple non sulfur bacteria B: 96 - sulfur bacteria B: 96 Pyrene, manganese nodules B: 125 Pyrite A: 107, 126 Pyro-phosphoric acid A: 147

Quartz A: 73 Quercus robus B: 5

Radiance geometry B: 138 -, spectral B: 139

313

314

Radiant energy B: 137, 138 - - of the surfaee B: 222 -, existanee B: 139 - exposure B: 138 - flux B: 139 - - density B: 236 Radiation B: 131, 209, 211, 220 - -, diural variations B: 221 - -, geographieal variations B: 221 - -, global mean B: 248 - -, seasonal variations B: 221 - -, top of the atmosphere B: 252 - dose, man B: 56 - effeets, natural radionuc1ides B: 56 - proeess, elementary B: 144 -, terrestrial B: 56 Radiative equilibrium B: 218 - heating eoo1ing B: 236 - transfer B: 200 - - equation B: 201 ff., 205 Radioearbon dating A: 101 Radionuc1ides, eosmogenie B: 50, 52 -, natural B: 47, 53 -, terrigenous B: 50 Radium, Isotopes B: 49 Rain A: 120 Rainforest latosol B: 34 Rainwater A: 41 Ratio of foreign broadening, water vapor

B: 163 Rattus norvegieus, generations B: 15 Rayleigh approximation B: 151 - optiea1 depth B: 177 - seattering B: 146, 148, 149, 177, 179 - - eoeffieient B: 173, 176 Rayleigh's theory B: 145 y-ray radiation B: 141, 143 Realized niehe B: 12 Redox potential, seawater A: 63 - reaetions in water A: 38 Redox-switeh meehanism A: 174 Reefs, primary produetion B: 25 Reflectance B: 142 - of c10uds B: 187 -, land surfaces B: 195 -, sand B: 196 - of urban, area B: 196 -, vegetated surfaces B: 196 - of water B: 192 Reflection B: 140 - funetion B: 142 - - of oeean water B: 190 - indicatrix B: 142 Refractive index B: 144, 147 - - of air B: 146 Relative airmass B: 173 Residence time for atmospheric lead A: 190

- -, lead A: 186 - -, mercury A: 180 Residual bands B: 194 "Reststrahlen" B: 194 Rhizobium B: 64

Subject Index

Roots, scavenging phosphate A: 157 Rossby waves B: 242

Sagebrush B: 35 Sahara, dust B: 265 -, - outbreaks B: 262 Sahel zone B: 278 Salinity A: 52 Salvia leucophylla B: 8 Saprotrophs B: 4 Savannah, net primary production B: 87 Savannah, primary production B: 25 Sawfly B: 18 Seale B: 133 - height of the atmosphere B: 176 Scattered radianee B: 202 - radiation B: 151 - solar radiation B: 179 Scattering B: 145 - eoefficient B: 140 - -, volume B: 149 - cross section B: 148 - efficiency B: 151 - phase function B: 146, 149, 202 Sehumann-Runge bands B: 164 Sehuster-Sehwarzsehild approximation B: 203 Sea breeze B: 243 - spray A: 42 - surface B: 190 - water, natural radionuc1ides B: 55 Seawater, buffering capacity A: 63 -, halogenated natural produets A: 250 -, properties A: 51, 53 Seeondary consumers B: 24 - productivity, ecosystem B: 25 Sediment accumulation rate A: 100 - - velocity A: 99 - percolation layer B: 137 -, phosphate A: 163 Sedimentary carbon A: 101 - organic matter A: 98 - reeord A: 10 - rocks A: 71 - sulfide A: 10 I - type B: 27 Sedimentation B: 103 - proeesses, isotopes B: 58 -, reeent aneient B: 112 Sedimentation, sulfur A: 122 Sediments A: 61 -, lead A: 188

Subject Index

Selective absorption B: 154 Selenium, methylation A: 217 Selinene A: 237 Semi-deserts B: 35 Senescence B: 14 Sensible heat flux B: 225, 229, 230 Shannon-Wiener index B: 22 Shores B: 38 Shortwave radiation in ocean water B: 192 Shrubs B: 20 Sierozem B: 34 -, grey desert soi! B: 35 Silt A: 72 Simpson's index B: 22 Single scattering albedo B: 202 Sky radiation B: 179 Smelting, lead emission A: 191 Snow albedo B: 197 SOz see sulfur dioxide Social ecology B: 2 Soil B: 6, 230 - air A: 81 -, chemical aspects A: 69 ff. - formation A: 82 -, major components A: 73 - organic constituents A: 80 - particles A: 75 - phosphate A: 153 - structure A: 82 - units A: 83 -, volume composition A: 72 - water A: 81 Soils A: 74 -, ecology B: 10 -, lead A: 188 -, tin A: 198 Solar atmosphere B: 171 - Constant B: 171, 212, 219, 220, 270 - flux, short-term fluctuations B: 213 - infrared radiation B: 141, 143 - irradiance B: 171, 177, 209, 212 - -, atmosphere B: 172 - radiance B: 171 - radiant energy B: 133 - radiation B: 249 - -, absorbed B: 251 - -, scattered B: 180 - spectral irradiance B: 171 - spectrum B 169 Solid waste, lead burden A: 193 Solubility equilibria A: 36 Sorption see adsorption, biosorption South Pole Station, trace elements B: 266 "Southern Oscillation" B: 246 Spatial niche B: 11 Species diversity B: 21 -, ecology B: 4

Specific growth rate B: 14 Spectrum B: 141

315

Spheric atmosphere, optical path B: 178 Sponges, organohalogen compounds A: 240 Spruce budworm, population growth B: 17 Squalene A: 238 Stability, ecological systems B: 30 Standard atmosphere B: 173 Standing crop B: 26 Stanols B: 115 Steady state B: 28 - -, ecological system B: 30 Stefan-Boltzmann constant B: 217 Stenohaline B: 9 Stenothermal B: 9 Steradienes B: 115 Sterenes B: 115 -, saturated B: 117 Steroids B: 115 - in geochemistry B: 117 Sterol patterns B: 117 Sterols in sediments B: 115 Stratification, ecology B: 20 Stratopause B: 136 Stratosphere 13 B: 136, 238, 239 Stratospheric aerosollayer B: 269 Strong line approximation B: 157 Structural ecology B: 2 - relationships, ecology B: 7 Stylatulide A: 245 Stylocheilamide A: 231 Subclimax B: 29 Succession, cyclic B: 28 -, ecology B: 28 -, primary B: 28 -, secondary B: 28 Sugar maple, nutrient quantities B: 7 Sulfate in groundwaters A: 113 - in the oceans A: 111 - reduction A: 135; B: 105 - in river water A: 113 Sulfate-reducing bacteria A: 93 Sulfide, autoxidation A: 127 - in marine system A: 112 Sulfur aerosols A: 108 -, annual fluxes A: 117 - compounds, biological oxidation A: 136 - -, biological reactions A: 132 - - in biological systems 128 - -, concentration in the atmosphere A: 107 - -, microbial production A: 138 - -, physical and chemical conversion A: 125 - cycle A: 105ff. - -, global A: 116 - dioxide A: 108; B: 160 - -, adsorption by the oceans A: 122 - -, oxidation A: 125

316

Sulfur emission A: 118 -, equilibrium chemistry A: 123 - hexafluoride in the atmosphere A: 107 -, inorganic compounds A: 124 - isotopes, fractionation by biological

processes A: 138 -, marine sediments A: 123 - in rain A: 120 -, residence time A: 107 -, world's reserves A: 114 -, methylation A: 217 Surface albedo feedback B: 270 - layer B: 136 - waters, lead A: 191 Surfaces, radiative properties B: 190 Surugatoxin A: 246 Survivorship curves B: 14 Swamp and marsh, net primary production

B: 87 B: 18

B: 286 Sympatry Symbols ß-synderol Synecology

A. 235 B: 126

Taiga podzol B: 34 Taigas B: 34 Tellurium, methylation A: 217 Temperate grassland, net primary production

B: 87 - shrubland B: 34 Temperature, absolute B: 132 - change due to constituent variations B: 258 - conditions, ecology B: 9 - increase B: 260 - inversion B: 238 -, potential B: 136 - profile of the atmosphere B: 238 - regions B: 22 Terrestrial communities B: 20 - ecosystem B: 23, 86 - emission B: 205 - - spectrum B: 182 - infrared radiation B: 141, 143 - material B: 126 - organisms, organohalogen compounds

A: 248 - system, partitioning ofthe global inventories of

nitrogen B: 72 - temperatures B: 217 Terrigenous radionuc1ides B: 50, 55 Tertiary consumers B: 24 Tetrabromopyrrole A: 230 Thallophytes B: 20 Thallium, biomethylation A: 215 The1epin A: 247 Thermal conductivity B: 223, 225 - diffusivity B: 224, 225

Subject Index

Thermoc1ine B: 137,243 Thermodynamic equilibrium B: 166 Thermosphere B: 136, 166, 239 Thorium decay series B: 51, 55 Thomwood B: 34 Three-spined stickleback B: II Thysiferol A: 238 Tidal zone B: 20 Time B: 213 Tin, biogeochemical cyc1es A: 197 -, biomethylation A: 201 -, environmentallevels A: 197 -, microbiological degradation A: 200 - stabilizers A: 199 Total organic carbon A: 40 Toxicities of elements A: 211 Trace elements in biochemistry A: 210 - gases B: 259 - metals, atmospheric vs fluvial transport

A:45 Tracers, geochemical B: 57 Trade wind B: 244, 246 - - regions B: 241 Trafik, lead levels A: 192 Transformation of sulfur in the environment

A: 125 Transmethylation A: 174 Transmission function B: 168,206 Transmittance B: 156 Transport see environmental transport - processes, natural radionuc1ide tracers B: 57 Trans-I,3,3-tribromo-I-heptene oxide A: 232 Transuranic elements B: 48 Travelling atmospheric systems B: 233 Trees B: 20 Trialkyl tin compounds, biocide A: 202 Triassic A: 14 Triediol A: 238 2,6, I 0-Trimethyl-dodecane B: 113 2,6,IO-Trimethyl tetradecane B: 113 Triphosphoric acid A: 147 Triterpenoids B: 118 - in geochemistry B: 123 Trophic niche B: 11 - structure, ecology B: 24 Tropical forest, primary production B: 25 - rain forest B: 279 - rainwoods B: 278 Tropopause B: 136 Troposphere A: 2; B: 134, 237 Tundra, net primary production B: 87 Tundras B: 35 Turbid atmosphere, radiation in B: 202-204 Turbidity factor B: 173-175, 178 Tumover, carbon dynamics B: 88 - time B: 27 Tyrian purple A: 247

Subject Index

Ultraviolett radiation B: 141, 143 Upwelling of oceans waters B: 244 Uranium decay series B: 55 -, isotopes B: 48 Urban atmosphere B: 179 UV radiation B: 164

Vagility B: 16 Van Allan belts B: 169 Vibration-rotation spectra B: 159 Violacene A: 233 Viscous sub-Iayer B: 137 Visible light B: 141, 143 Vitamin B12 A: 175 Voigt line profile B: 155 Volcanic eruptions B: 264, 265, 269 - events B: 263 - gases A: 6, 94 - rocks A: 71 - sulfur A: 118 Volcanoes B: 267 Von Korman constant B: 227 "Vorticity" B: 242

Walker circulation B: 246 Warm water sphere B: 137, 207 Water see also seawater B: 184, 228, 246 -, acidity A: 29 -, chemistry of natural A: 26 - clouds B: 187 - content of clouds B: 185 -, ecology B: 9 -, mercury content A: 181

-, molecule A: 18 -, natural composition A: 44 -, physical properties A: 21 -, properties A: 17 -, redox chemistry A: 38 -, reflectance spectrurn B: 193 -, reservoirs A: 47 -, structure of liquid A: 20 - surfaces, optical properties B: 190 - vapor B: 162,239,259 - - absorption B: 159 - - mass B: 174 - -, rotation spectrum B: 158 - -, stratospheric B: 259 - - transport B: 228, 251 -, viscosity A: 25 Wave breaking layer B: 136 Waves in the thermosphere B: 240 Weathering A: 69, 82; B: 103 - reactions A: 42 West wind regime B: 244 Whiting B: 98 Wind profile B: 226 Woodlands B: 34

317

World Climate Research Programme B: 238 - Weather Watch B: 283 Worms A: 247

X-ray radiation B: 141, 143

Zenith radiance B: 179 Zone Mean Times B: 214

The Handbook of Environmental Chemistry Editor: O. Hutzinger This handbook is the first advanced level compendium of environmental chemistry to appear to date. It covers the chemistry and physical behavior of compounds in the environment Under the editorship ofPro( O. Hutzinger, directoroftheLaboratory ofEnvironmentai and ToxicologicaI Chemistry at the University of Amsterdam. 37 international specialists have contributed to the first three volumes. For a rapid publication ofthe material each volume is divided into two parts:Each volume contains a subject index.

Tbe Handbook of Environmental Chemistry is a critical and complete outline of our present knowledge in this field and will prove invaluable to environmental scientists, biologists, chemists (biochemists, agricultural and analytical che~ists), medical scientists, occupationaI and envIronmenta! hygienists, research geologists, and meteorologists, and industry and administrative bodies.

Springer-Verlag Berlin Heidelberg NewYork

Volume 1 (in 2 parts) PartA

The Natural Environment and the Biogeoehemieal Cycles With contributions by numerous experts 1980. 54 figures. XV, 258 pages ISBN 3-540-09688-4

Contents: The Atrnosphere. - The Hydrosphere. - ChemicaI Oceanography. - Chemical Aspects ofSoiI. - The Oxygen Cyde. - The Sulfur Cyde. - The Phosphorus Cyde. - Meta! Cydes and Biological Methylation. - Natural Organohalogen Compounds. -Subject Index.


Reaetions and Proeesses With contributions by numerous experts 1980. 66 figures, 27 tables. XVIII, 307 pages ISBN 3-540-09689-2

Contents: Transport and Transfonnation of Chemicals: A Perspective. - Transport Processes in Air. - Solubility, Partition Coefficients, Volatility, and Evaporation Rates. - AdsOiption Processes in SoiI. - Sedimentation Processes in the Sea. - Chemical and Photo Oxidation. - AtmosphericPhotochemistry.Photochemistry at Surfaces and Interphases. -Microbial Metabolism. - Plant Uptake, Transport and Metabolism. - Metabolism and Distribution by Aquatic Animals. -LaboratoIY Microecosystems. - Reaction Types in the Environment. -Subject Index.


Anthropogenie Compounds With contributions by numerous experts 1980.61 figures, 73 tables. XIII, 274 pages ISBN 3-540-09690-6

Contents: Mercury. - Cadmium. - Polycydic Aromatic and Heteroaromatic Hydrocarbons. - Fluorocarbons. -Chlorinated Paraffins. - Chloroaromatic Compounds Containing Oxygen. - Organic Dyes and Pigments. - Inorganic Pigments. - Radioactive Substances. - Subject Index.

Springer-Verlag Berlin Heidelberg NewYork

L. G. NickeIl

Plant Growth Regulators Agricultural Uses

1982. 29 figures. XII, 173 pages ISBN 3-540-10973-0

Plant growth regulators can be defined as either natural or syntheticcompounds thatare applied directly 10 a target plant to alter its life processes or its structure 10 improve quality, increase yields, or facilitate harvesting. The author ofthis book - Vice President ofResearch and Development at the Velsicol Chemical Corporation, Chicago, Chairman ofthe Plant Growth Regulator Working Group, and the Treasurer ofthe American Society ofPlant Physiologists - discusses the effects ofthese regulators on plant functions such as rootinduction, control offlowering, control of sex, and control of maturation and aging. Emphasis is placed on the practical aspects of growth regulators rather than their mode of action. An extensive bibliography is provided which inc1udes reviews by specialists in all areas of plant growth and development, both basic and applied. Special consideration is given to the current commercial uses of plant growth regulators for a variety of purposes on a number of economicaHy important crops as weH as to the direct and indirect financial returns to the grower. Tables are provided which showthe chemical name, common name, trademark identification, code designation, producer companies, as weH as plant growth regulatory activity and, where applicable, other biological activities of the regulators discussed in the text. The book will be of interest to teachers, students and specialists in agronomy horticulture, plant physiology, crop science, nursery science, landscape architecture, pomology, seed technology, providing them with the only modern treatment of the uses of chemicals for the control of plant growth. (1201 ref.)

C.Pedtke

Biochemistry and Physiology of Herbicide Action 1982. 43 figures, 58 tables. XI, 202 pages ISBN 3-540-11231-6

the natural environment and the biogeochemical cycles

Documents