rad forc nat aer review
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Atmospheric Environment 39 (2005) 20892110
Radiative effects of natural aerosols: A review
S.K. Satheesha,, K. Krishna Moorthyb
aCentre for Atmospheric and Oceanic Sciences, Indian Institute of Science, Bangalore-560 012, IndiabSpace Physics Laboratory, Indian Space Research Organisation (ISRO), Vikram Sarabhai Space Centre, Trivandrum-695 022, India
Received 22 March 2004; received in revised form 30 September 2004; accepted 9 December 2004
Abstract
In recent years, there has been a substantial increase in interest in the influence of anthropogenic aerosols on climate
through both direct and indirect effects. Several extensive investigations and coordinated field campaigns have been
carried out to assess the impact of anthropogenic aerosols on climate. However, there are far fewer studies on natural
aerosols than on anthropogenic aerosols, despite their importance. Natural aerosols are particularly important because
they provide a kind of base level to aerosol impact, and there is no effective control on them, unlike their anthropogenic
counterparts. Besides, on a global scale the abundance of natural aerosols is several times greater than that of the major
anthropogenic aerosols (sulphate, soot and organics). The major natural aerosol components are sea salt, soil dust,
natural sulphates, volcanic aerosols, and those generated by natural forest fires. As with anthropogenic aerosols, the
abundance of natural aerosols such as soil dust is also increasing, due to processes such as deforestation, which exposes
more land areas which may then interact directly with the atmosphere, and due to other human activities. Since a major
fraction of the natural aerosol (sea salt and natural sulphate) is of the non-absorbing type (and hygroscopic), it partly
offsets the warming due to greenhouse gases as well as that due to absorbing aerosols (e.g., soot). The mineral dusttransported over land and ocean causes surface cooling (due to scattering and absorption) simultaneously with lower
atmospheric heating (due to absorption); this could in turn intensify a low-level inversion and increase atmospheric
stability and reduce convection. To accurately predict the impact of dust aerosols on climate, the spatial and temporal
distribution of dust is essential. The regional characteristics of dust source function are poorly understood due to the
lack of an adequate database. The reduction of solar radiation at the surface would lead to a reduction in the sensible
heat flux and all these will lead to perturbations in the regional and global climate. Enhanced concentration of sea salt
aerosols at high wind speed would lead to more condensation nuclei, increase in the cloud droplet concentration and
hence cloud albedo. Even though direct radiative impacts due to sea salt and natural sulphate are small compared to
those due to anthropogenic counterparts, their indirect effects (and the uncertainties) are much larger. There is a
considerable uncertainty in sea salt aerosol radiative forcing due to an inadequate database over oceans. The presence
of natural aerosols may influence the radiative impact of anthropogenic aerosols, and it is difficult to separate the
natural and anthropogenic aerosol contributions to radiative forcing when they are in a mixed state. Hence it isnecessary to document the radiative effects of natural aerosols, especially in the tropics where the natural sources are
strong. This is the subject matter of this review.
r 2005 Elsevier Ltd. All rights reserved.
Keywords: Aerosols; Climate change; Radiative forcing; Radiation budget
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www.elsevier.com/locate/atmosenv
1352-2310/$ - see front matterr 2005 Elsevier Ltd. All rights reserved.
doi:10.1016/j.atmosenv.2004.12.029
Corresponding author. Tel.: +91 80 22933070/22932505; fax: +91 80 23600865.
E-mail address: [email protected] (S.K. Satheesh).
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1. Introduction
The global climate system is a consequence of
interactions between its sub-components (Shaw, 1983;
Murthy, 1988; Charlson et al., 1992; Andreae, 1995;
Hansen et al., 1997, 1998; Clarke, 1998; Haywood et al.,
1999; Prospero et al., 2002; Seinfeld et al., 2004). The
main processes that determine the overall state of the
climate system are heating by incoming solar radiation
and cooling by outgoing long-wave (infrared) terrestrial
radiation (Coakley et al., 1983; Ramanathan et al., 1989;
Charlson et al., 1991; Kiehl and Briegleb, 1993; Hansen
et al., 1998; Seinfeld et al., 2004). Any process that can
disturb the overall energy balance can cause climate
change or perturbation (Kaufman and Fraser, 1997;
Kaufman et al., 1997; Seinfeld and Pandis, 1998). A
process that alters the radiative balance of the climate
system is known as radiative forcing (Coakley et al.,
1983; Coakley and Cess, 1985; Ramanathan et al., 1989;Charlson et al., 1991, 1992; Hansen et al., 1997, 1998;
Russell et al., 1999; Bates, 1999; Raes et al., 2000).
Radiative forcing can be internal or external. External
forcing operates from outside the Earths climate system
and includes orbital variations and changes in incident
solar flux. Volcanic activity is an example of an internal
forcing mechanism (Hoffmann et al., 1987; Moorthy
et al., 1996; Soden et al., 2002). Similarly, changes in the
composition of the atmosphere constitute another major
internal forcing mechanism, and the best examples are
the greenhouse gases and aerosols (Shaw, 1983; Crutzen
and Andreae, 1990; Charlson et al., 1992; Clarke, 1993;
Kaufman et al., 1997; Bates, 1999; Bates et al., 2000;
Rodhe, 2000; Prospero et al., 2002). Changes in the
greenhouse gas or aerosol content of the atmosphere
affects the radiative balance of the climate system
(Haywood and Ramaswamy, 1998; Myhre et al., 1998;
Haywood et al., 2003).
The Earths climate is strongly influenced by the
manner in which solar radiation is absorbed and
reflected in the atmosphere (Chylek and Wong, 1995;
Schwartz et al., 1995). During the past 100 years the
amount of carbon dioxide in the atmosphere has
increased by about 25% on account of human activities
(fossil fuel/biomass burning) (Meehl et al., 1996;Le Treut et al., 1998; IPCC, 2001). This has caused
the surface temperature of the Earth to increase globally
by about one kelvin (Meehl et al., 1996; Le Treut
et al., 1998).
In recent years, there has been a substantial increase
in interest in the influence of anthropogenic aerosols on
the climate through both direct and indirect radiative
effects. Several extensive investigations and coordinated
field campaigns have been carried out to assess the
impact of anthropogenic aerosols on climate. However,
studies of natural aerosols are few compared to those of
anthropogenic aerosols, despite the importance of the
former. Among these studies, Aerosol characterization
experiment-1 (ACE-1) focussed on natural aerosols.
Aerosol characterization experiments (ACE) were de-
signed to increase the understanding of how atmo-
spheric aerosol particles affect the Earths climate
system (Bates, 1999; Russell and Heintzenberg, 2000;
Seinfeld et al., 2004). ACE-1 was conducted over
southern hemispheric mid-latitudes with a specific goal
of understanding the properties and controlling factors
of aerosols in the remote marine atmosphere that are
relevant to radiation balance and climate (Bates, 1999;
Hainsworth et al., 1998; Griffiths et al., 1999). This
environment provided an opportunity to establish the
chemical, physical and radiative properties of a natural
aerosol system. ACE-2 was conducted during July 1997
to study the radiative effects of anthropogenic aerosols
from Europe and desert dust from Africa as they are
transported over the North Atlantic Ocean (Russell and
Heintzenberg, 2000). While ACE-1 and ACE-2 focussedmostly on natural and anthropogenic aerosols, respec-
tively, ACE-Asia focussed on a complex mix of
anthropogenic and natural aerosols over the Asian
region (Huebert et al., 2003; Seinfeld et al., 2004).
In this paper, we review the role of natural aerosols in
modifying the Earths radiation budget and demonstrate
its importance in the climate change debate. Throughout
this paper we use the term aerosols to address to the
particulate phase of the atmospheric aerosol system.
2. Earths radiation balance: role of aerosols
The Suns radiation, much of it in the visible region of
the spectrum, warms our planet. On average, the Earth
must radiate back to space the same amount of energy
that it gets from the Sun (Seinfeld et al., 2004).
Greenhouse gases (GHGs) in the Earths atmosphere,
while largely transparent to incoming solar radiation,
absorb most of the infrared (IR) radiation emitted by
the Earths surface. Clouds also absorb in the IR. Thus,
part of the IR emitted by the surface gets trapped (and
this is the natural greenhouse effect). Under a clear sky,
about 6070% of the natural greenhouse effect is due to
atmospheric water vapour (Seinfeld and Pandis, 1998).The next most important GHG is carbon dioxide,
followed by methane, ozone, and nitrous oxide.
If we represent solar radiation incident at the top of
the atmosphere (global) as 100 units, then a net amount
of 51 units reaches the surface (Fig. 1). Of the remaining
49 units, 3 units are absorbed by clouds and 16 units by
aerosols, water vapour and CO2 together. The clouds,
surface and atmosphere (which include aerosols as well)
reflect 17 units, 6 units and 7 units, respectively. Of the
51 units absorbed by the Earths surface, 23 units are
released as latent heat, 7 units as sensible heat and 21
units as infrared. About 15 units of infrared are
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absorbed by aerosols, water vapour and CO2. Thus,
aerosols have an important influence on the Earths
radiation balance.
It is widely known that warming, which tends to
enhance evaporation, will increase the water vapour
content of the troposphere (Seinfeld et al., 2004). This
further amplifies the warming, as water vapour is a
dominant GHG. Snow and ice reflect much of the
incident sunlight back to space; thus a reduction of snow
and ice cover would also lead to enhanced warming.
Clouds are generally good absorbers of infrared, but
high clouds have cooler tops than low clouds, so they
emit less infrared spaceward. The interplay between
atmosphere (GHGs, molecular absorbers and aerosols),
ocean, clouds, and ice is poorly understood.
The mean temperature of the Earth (Te) is given by
the balance between the absorbed solar energy and
emitted terrestrial energy given by the steady state
condition,
H S0
41 a sT4e 0, (1)
where H is the net energy input to the climate system
and S0 is the solar power per unit area intercepted at themean SunEarth distance (solar constant)
(13651372 W m2) (Seinfeld and Pandis, 1998). The
factor 4 is the ratio of the Earths surface area to the
cross-sectional area. The quantity a is the albedo
(reflectance) of the Earth, which is the fraction of the
incident solar radiation reflected by the Earths surface
and atmosphere and has a mean value of $0.3.
Consequently, of the 343 W m2 of the mean solar
radiation incident at the top of the atmosphere,
$103Wm2 is reflected back to space by the Earths
surface and atmosphere. Aerosols can influence the
albedo and thus can have an impact on the climate
system. The energy balance equation implies that a
change of 0.01 in the value of a results in about 1%
change in the global temperature.
The question of whether aerosols increase or decrease
the value ofa (warm or cool the planet) depends on their
chemical composition. Completely scattering aerosols
will increase a (which means a decrease in temperature)
whereas absorbing aerosols (e.g., soot) would lead to a
decrease in a (Coakley and Cess, 1985; Hansen et al.,
1998; Russell et al., 1999) and hence an increase in Te.
This means the warming or cooling effect can change
from region to region depending on many factors such
as the relative strengths of various aerosol sources and
sinks (Kaufman et al., 1997; Clarke, 1998; Russell et al.,
1999; Ginoux et al., 2001, 2004; Ramanathan et al.,
2001; Luo et al., 2003). When the net effect of aerosols is
cooling, they partly offset the greenhouse warming,
while if the net effect is warming, they complement the
greenhouse warming (IPCC, 2001). Since aerosolproperties show large regional variations, the regional
impact can be very different, and this is the main reason
why the importance of aerosols is poorly characterised
in climate models. This is especially true for natural
aerosols, because of the lack of a comprehensive
database.
3. Radiative effects of natural aerosols
It is well known that aerosols are of natural or
anthropogenic origin. The source strengths of various
natural and anthropogenic species are given in Table 1
(data from Andreae, 1995). It can be seen that, in terms
of emission, natural aerosols contribute 89%. In terms
of column mass and optical depth, natural aerosols
contribute 81 and 52%, respectively. Thus there exists
no direct relationship between aerosol mass, optical
depth and its radiative impact. Out of the major natural
and anthropogenic aerosol types (sulphate, nitrate, sea
salt, carbonaceous matter (organic carbon and black
carbon), mineral dust, oceanic sulphate and so on) sea
salt, soil dust and oceanic sulphate constitute a major
portion of the global natural aerosol abundance (during
volcanically quiescent periods) even though a propor-tion of the dust could also be due to anthropogenic
activities (Tegen and Fung, 1994; Sokolik and Toon,
1999; Sokolik et al., 1998; Tanre et al., 2003; Haywood
et al., 2003; Highwood et al., 2003). So the accurate
estimate of natural/anthropogenic fraction is difficult to
determine. Another natural component of aerosol is
naturally occurring soot (smoke from natural burning
such as forest fires). Natural and anthropogenic soot is
the main absorbing fraction of aerosol (Crutzen and
Andreae, 1990; Chylek and Wong, 1995; Kaufman et al.,
1998; Jacobson, 2001; Babu and Moorthy, 2002; Sato
et al., 2003) and is among the most complex aerosol
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Fig. 1. Earths radiation budget demonstrating the role of
aerosols.
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types. It is produced both by natural and anthropogenic
processes such as forest fires, man-made burning or
combustion and transportation (Schwartz et al., 1995).
Its radiative effects vary depending on the production
mechanism. Soot has a significant role in climate
modification because of its absorption characteristics
(Haywood and Shine, 1995; Kaufman et al., 1997;
Haywood and Ramaswamy, 1998; Haywood and
Boucher, 2000; Babu and Moorthy, 2002). Even though
laboratory analysis can distinguish soot from biomassburning from that of fossil fuel origin, in a global
scenario it is not possible to quantify the natural fraction
of soot, and it is generally believed that a major fraction
of soot is produced by anthropogenic activities. Thus,
we focus more on sea salt, dust and oceanic sulphates.
However, for the purpose of comparison, we discuss
anthropogenic counterparts as well.
A simplified block diagram in Fig. 2 shows the
radiative effects of the three major natural aerosols
considered here. Although, the generation of sea salt
and dust depends primarily on the surface wind speed,
their subsequent upward transport depends on the
boundary layer characteristics, including mixing height,vertical winds and so on. These would be different over
land and sea. After production, dust aerosols are often
transported long distances from their sources (Arimoto
et al., 2001). Examples are dust transport from the
Sahara across the Atlantic Ocean, Arabian dust trans-
port across the Arabian Sea and dust from China across
the Pacific. Mineral dust is believed to play an important
role in marine biological processes (Falkowski et al.,
1998). For example, dust is a source of iron, which acts
as a nutrient for phytoplankton (Falkowski et al., 1998;
Fung et al., 2000). This, in turn, would influence
dimethyl sulphide (DMS) emission from the oceanic
phytoplankton and hence natural production of sul-
phate aerosols over the ocean. Natural sulphate aerosols
over oceans are good condensation nuclei for formation
of clouds. Charlson et al. (1987) hypothesised that there
exists a negative feedback mechanism by which an
increased number of natural sulphate aerosols over
oceans increases the cloud albedo and hence causes a
reduction of surface-reaching solar radiation. This, in
turn, reduces the DMS emission leading to a reduction
in the natural sulphate production rate (Fig. 2). This
hypothesis was extensively studied in experiments such
as ACE-1 and ACE-2. Similarly, sea salt aerosols are
also hygroscopic in nature and act as condensationnuclei for the formation of clouds.
Dust aerosols reduce the surface-reaching solar
radiation (due to scattering and absorption) while
heating the lower atmosphere (due to absorption). This
modifies the atmospheric boundary layer characteristics
over land and ocean (Fig. 2). Over the ocean an
increased concentration of dust also contributes to a
reduction of surface-reaching solar radiation. The
combined effect of these three major natural aerosols
may have an influence on sea surface temperature.
Detailed discussions on the radiative effects of each of
these aerosols are included in the following sections.
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Table 1
Source strength (data from dAlmeida et al., 1991; Andreae,
1995)
Source Emission,
Tgyr1Column
burden,
mg m2
Optical
depth
Natural
Primary
Soil dust 1500 32.2 0.023
Sea-salt 1300 7.0 0.003
Volcanic dust 33 0.7 0.001
Biological debris 50 1.1 0.002
Secondary
Sulphates 102 2.7 0.014
Organic matter 55 2.1 0.011
Nitrates 22 0.5 0.001
Total Natural 3060 46 0.055
Anthropogenic
Primary
Industrial dust 100 2.1 0.004
Black carbon 20 0.6 0.006
Secondary
Sulphates 140 3.8 0.019
Biomass burning (w/o BC) 80 3.4 0.017
Nitrates 36 0.8 0.002
Organic matter 10 0.4 0.002
Total Anthropogenic 390 11.1 0.050
Total 3450 57 0.105
Anthropogenic fraction 11% 19% 48%
Iron fertilisation dueto transported dust
Surface Winds
sea-salt Production
Rate
Reduction in Surface
Solar Flux over
Ocean
Cloud Cover
over Ocean
DMS Emission
over Ocean
Dust Production
Rate
Boundary
Layer
Properties
over Land
Reduction in
Surface
Solar Flux & Lower
Atmosphere
Heating
Cloud Formation
SSTCloud
Formation
Fig. 2. Block diagram showing the climate impact of natural
aerosols.
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The radiative effects of aerosols depend strongly on
the single scattering albedo (SSA) (ratio of scattering to
extinction), which in turn depends on the real andimaginary components of the aerosol refractive index.
The single scattering albedos of three major aerosol
species (sea salt, soot and dust) are shown in Fig. 3
(data from Hess et al., 1998). It may be noted that SSAs
of dust reported by various studies show discrepancies.
Despite these, it can be seen that in the visible region, sea
salt and sulphate are non-absorbing (SSA is close to
unity) and soot is highly absorbing (SSA is $0.23 at
500 nm). In the long-wave region, sea salt and sulphate
are partly absorbing and soot is completely absorbing.
Recent studies have shown that a proportion of the
mineral dust in the atmosphere may be of anthropogenic
origin and exert significant radiative forcing (Cattrall
et al., 2003; Haywood et al., 2003; Highwood et al.,
2003; Tanre et al., 2003). However, the optical and
radiative properties of dust are not known precisely. The
SSA of dust at 0.55 mm reported by Hess et al. (1998)
based on observations in the past was $0.84 (Fig. 3).
Recent studies have shown that the refractive indices
used for dust aerosols in global models are in error
(Kaufman et al., 2001; Haywood et al., 2003). Kaufman
et al. (2001), using remote sensing, inferred the SSA of
Saharan dust as 0.97 at 0.55 mm. The studies on Saharan
dust by Haywood et al. (2003) have shown that the SSA
of dust at 0.55mm is in the range of 0.950.99. Thesignificantly lower SSA reported in the past (Hess et al.,
1998, for example) could be due to the possible mixing
of Saharan dust with biomass (possibly soot) aerosols.
The Saharan dust experiment (SHADE) was designed to
better understand the controlling factors that determine
radiative forcing of dust (Haywood et al., 2003; Tanre et
al., 2003). These studies suggest that mineral dust has a
cooling effect and the model estimate of direct radiative
forcing of Saharan dust is 0.4Wm2 (Tanre et al.,
2003). In the terrestrial region, Saharan dust decreased
the upwelling radiation at the top of the atmosphere by
6.5Wm
2
and increased the surface radiation by
11.5Wm2 (Highwood et al., 2003). The SSA and
phase function of African mineral dust were retrieved at
14 wavelengths across the visible spectrum from ground-based measurements (Cattrall et al., 2003). The SSA
showed a spectral shape expected of iron-bearing
minerals but is much higher than climate models have
assumed, indicating that wind-blown mineral dust cools
the Earth more than is generally believed (Haywood
et al., 2001; Catrall et al., 2003).
The radiative effects of aerosols depend on the type
and altitude of clouds as well (Heintzenberg et al., 1997;
Satheesh, 2002a, b). In Fig. 4, we show a representation
of a cloudy atmosphere. In case (a) most of the aerosols
are concentrated below clouds whereas in case (b)
aerosols are mostly above clouds. The radiative impact
of aerosols in case (a) and case (b) can be significantly
different even when aerosol column properties are the
same. When a cloud layer is present above aerosols,
most of the incident radiation will be reflected back and
a small fraction only will interact with aerosols. On the
other hand when an elevated aerosol layer is present
with a cloud below, the aerosols interact not only with
radiation incident from the Sun, but also with that
reflected from the cloud layer below. This would result
in an enhanced aerosol radiative impact.
3.1. Sea salt aerosols
The strongest natural aerosol production rate is that
of sea salt, at an estimated 100010,000 Tg per year
(Winter and Chylek, 1997). This is about 3075% of all
natural aerosols (Blanchard and Woodcock, 1980). The
source of airborne salt particles is obviously the sea. But
most of the early investigators did not concentrate on
the exact mechanism of production of sea salt particles.
In the light of laboratory experiments, Stuhlman (1932)
reported that the bursting of bubbles in distilled water
produced jets of water which broke into small droplets.
Later, Kohler (1936, 1941) proposed that the formation
of spray at the wave crest by strong winds was
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0.00
0.20
0.40
0.60
0.80
1.00
0 5 10 15 20 25 30 35 40
Wavelength (m)
SingleSca
tteringAlbedo
Dust Soot Sea-salt
Fig. 3. Single scattering albedo for major aerosol species.Fig. 4. The effect of clouds on aerosol radiative forcing.
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responsible for the airborne salt particles. The high-
speed photographic study of Kientzler et al. (1954) of
bursting bubbles confirmed the mechanism of ejection of
droplets from a breaking bubble as suggested by
Stuhlman.
A number of investigators have studied the wind
speed dependence of the concentration of sea salt
particles in the marine boundary layer (Woodcock,
1953, 1957; Monahan, 1968; Blanchard and Syzdeck,
1988; Tsunogai et al., 1972; Lovett, 1978; Monahan
et al., 1982, 1983; Hoppel et al., 1990; Parameswaran et
al., 1995; Moorthy et al., 1997; Moorthy and Satheesh,
2000; Vinoj and Satheesh, 2003). These showed a clear
dependence of sea salt aerosol mass concentration on
wind speed. Many of these investigators have suggested
an exponential relation of the form
C C0 expbU, (2)
where C is the aerosol number or mass concentration atwind speed U, C0 that at U 0 and b is a wind index.
There have been a few experiments to understand the
effects of natural aerosols on climate. Among these,
ACE-1 was one of the major experiments. ACE-1
focussed on remote marine aerosol, minimally perturbed
by continental sources, whereas ACE-2 studied the
outflow of European aerosol into the northeast Atlantic
atmosphere (Bates et al., 2000; Quinn et al., 2000).
During ACE-2 sub-micrometre aerosol dominated
scattering by the whole aerosol in contrast to ACE-1
where super-micrometre aerosol was the dominant
scatterer. During the first aerosol characterisation
experiment (ACE-1), extensive studies were carried out
on the influence of sea salt on aerosol radiative
properties (Murphy et al., 1998). However, in both
ACE-1 and ACE-2, there was poor correlation between
local wind speed and sea salt mass concentration (Quinn
et al., 2000). On the other hand, many investigations
have observed a correlation between aerosol character-
istics and averaged wind (ODowd and Smith, 1993;
Parameswaran et al., 1995; Moorthy et al., 1997, 1998;
Satheesh et al., 2002). It may be noted that there is a
possibility of sea salt advection from regions of high
wind to regions where wind speeds are low ( Gong et al.,
1997, 2002; Kinne et al., 2003) and this can result in ahigh aerosol load even over regions of low winds. This
might possibly explain the observations of poor
correlation between local wind speed and sea salt mass
concentration during ACE (Quinn et al., 2000).
Detailed estimates of sea salt aerosol radiative forcing
(Winter and Chylek, 1997) showed that at low wind
speed, the sea salt radiative forcing is in the range of
0.6 to 2 W m2 and at higher wind speeds this can be
as high as 1.5 to 4 W m2. This negative forcing by
naturally occurring sea salt aerosol is quite significant
when we consider the fact that forcing caused by
projected doubling of CO2 is about +4 W m
2
. The
forcing caused by the increase in CO2 since the advent of
the industrial era is about +1.46 W m2 (Charlson et al.,
1992; Winter and Chylek, 1997). It may be noted that
there are very few data on sea salt aerosols where wind
speeds are high. In such conditions the measurements
are difficult. Thus there is a considerable uncertainty in
sea salt aerosol radiative forcing (Gong et al., 2002;
Kinne et al., 2003).
Another recent study has demonstrated that as wind
speed increases there are two competing effects which
determine the aerosol forcing at the surface; they are the
increase in the single scattering albedo (SSA) and the
increase in the optical depth (Satheesh, 2002a, b;
Satheesh and Lubin, 2003). An increase in single
scattering albedo decreases the forcing efficiency at the
surface whereas an increase in optical depth increases
the forcing (Heintzenberg et al., 1997). But at the top of
the atmosphere (TOA), increases in both SSA and
optical depth increase the forcing. The study has shownthat as the sea-surface wind speed increases from 0 to
1 5 m s1, the magnitude of aerosol forcing at the TOA is
enhanced by $6 W m2 (i.e., larger negative value)
(Satheesh, 2002a, b; Satheesh and Lubin, 2003). It may
be noted that the magnitude of composite aerosol
forcing at the TOA observed over the tropical Indian
Ocean was only $1 0 W m2 (Satheesh and Rama-
nathan, 2000; Satheesh et al., 2002). This shows that
modulation in forcing by sea salt aerosols (produced
by sea-surface winds) is quite significant. It also
demonstrates that surface wind has a significant role in
changing the chemical composition of aerosols over the
sea and hence the forcing (Satheesh and Lubin, 2003).
Aerosol short-wave, long-wave and net forcing as a
function of wind speed is shown in Fig. 5 (data from
Satheesh and Lubin, 2003). These values are in
agreement with those reported by Winter and Chylek
(1997). Model estimates of aerosol forcing in clear and
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-8.5
-7.5
-6.5
-5.5
-4.5
-3.5
-2.5
-1.5
-0.5
0.5
1.5
2.5
2 4 6 8 10 12
Wind Speed (m s-1)
AerosolTOAForcing(Wm
-2)
SW LW Net
Fig. 5. Aerosol forcing as a function of wind speed (based on
data reported in Satheesh and Lubin, 2003).
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cloudy skies (when absorbing aerosols are present) have
shown that aerosol forcing at the TOA decreases as
cloud cover increases and can be positive if cloud
coverage exceeds $25% (Podgorny and Ramanthan,
2001; Satheesh, 2002a, b). When a reflecting cloud layer
is present (see Fig. 4), both aerosol scattering and
absorption effects are amplified due to the multiple
interactions of the radiation reflected back by clouds or
between clouds and surface (Heintzenberg et al., 1997;
Satheesh, 2002a, b). The effect of the sea-surface winds is
thus to offset part of the heating by absorbing aerosols
(as the TOA forcing by sea salt aerosol is negative in
both clear and cloudy skies).
3.2. Oceanic sulphate aerosols
There is evidence that fine particles are produced over
the sea (Clarke et al., 1987, 1997; Clarke, 1993; Hoppel
et al., 1990; Fitzgerald, 1991; Hoppel et al., 1994; Pandiset al., 1994; Russell et al., 1994; Bates et al., 2000; Quinn
et al., 2000; Johnson et al., 2000; Putaud et al., 2000;
Clarke and Kapustin, 2002). The resulting particles,
after subsequent growth by condensation and coagula-
tion to larger sizes (radius, R40.1mm), play a dominant
role in producing the marine stratocumulus clouds by
acting as cloud condensation nuclei (CCN) over remote
oceanic regions (Hoppel et al., 1986; Clarke, 1993;
Lawrence, 1993; Russell et al., 1994; Bates et al., 2000;
Johnson et al., 2000; Putaud et al., 2000). Aerosol
measurements made over the tropical oceans have
shown that the sub-micrometre aerosol size distributions
can be constant for a week or longer irrespective of the
prevailing meteorological conditions (Clarke et al., 1987,
1997; Hoppel et al., 1986, 1990; Pandis et al., 1994).
Several investigations in clean marine air have shown
that most of the particles o0.25mm are composed of
non-sea salt sulphate. Aerosol volatility measurements
are in good agreement with this fact (Clarke et al., 1987;
Fitzgerald, 1991; Clarke, 1993; Pandis et al., 1994). The
studies as part of ACE-1 have shown that new particles
are not formed in abundance in the marine boundary
layer, but rather in the relatively particle-free atmo-
sphere of the upper troposphere (at least above the
marine boundary layer) (Bates et al., 2000; Quinn et al.,2000). Particles from gas-to-particle conversion are more
volatile than sea salt and can be distinguished from sea
salt by measuring the temperature at which the particles
decomposed (Fitzgerald, 1991).
By measuring the volatility of particles, Clarke et al.
(1987) have shown that approximately 99% of the
particles smaller than 0.2 mm radius behaved like
sulphuric acid or ammonium sulphate/bisulphate and
particles with r40.25 mm behaved like sea salt. Hoppel
et al. (1990) measured the volatility of sub-micrometre
particles (ro0.3mm) over remote parts of the Pacific
Ocean and found that most of the particles were non-sea
salt particles except during stormy periods, during which
enough salt particles can be produced. It is believed that
a significant fraction of the tropospheric aerosol mass
over oceans in the sub-micrometre size range is
principally derived from homogeneous in-cloud oxida-
tion of gaseous sulphur compounds (Charlson et al.,
1987; Langner et al., 1992; Clarke, 1993). The sulphur
compounds present over remote oceans can be of marine
or continental origin.
Consideration of the source strengths of various
organo-sulphur gases emitted by the ocean and their
rate constants for oxidation by the hydroxyl ion have
lead to the conclusion that DMS is the major source of
non-sea salt sulphate over oceans (Andreae et al., 1983;
Fitzgerald, 1991). Charlson et al. (1987) have also
proposed that DMS is the major source of aerosol
sulphate in the remote marine atmosphere. Natural
emissions of sulphur represent a significant part of the
total flux of gaseous sulphur to the atmosphere(Andreae, 1985). Almost all species of marine phyto-
plankton release DMS as DMS vapour, which gets
oxidised by different radicals to form SO2 (Fitzgerald,
1991; Russell et al., 1994). In the atmosphere DMS is
oxidised by the several radicals including OH, NO3 and
IO (Fitzgerald, 1991), the OH radical being the major
oxidant. Photo-oxidation of DMS (CH3SCH3) with
OH yields SO2, methane sulphonic acid (MSA), H2SO4and numerous other compounds (Russell et al., 1994;
Fitzgerald, 1991). The photo-oxidation products of
DMS are converted to non-sea salt sulphate by gas-to-
particle conversion processes (Fitzgerald, 1991). These
non-sea salt sulphate particles grow by acid condensa-
tion to a radius of$0.04 mm in about two days where the
particles are large enough to act as CCN, and can grow
further while cycling through non-precipitating clouds
(Hoppel et al., 1994). The non-sea salt sulphate particles
present in the marine atmospheric boundary layer
(MABL) play an important role in acting as CCN
(Charlson et al., 1987; Lawrence, 1993; Clarke, 1993).
The number of these particles capable of acting as CCN
varies from $30 to 200cm3 (Pruppacher and Klett,
1980; Clarke et al., 1987; Hoppel et al., 1990, 1994).
Sulphate aerosols present over oceans can be of
natural or anthropogenic origin. Though anthropogenicsulphur emissions can influence the sulphate concentra-
tion over oceans, in most of the remote areas of oceans,
natural emissions of sulphur can account for almost all
non-sea salt sulphate (Savoie and Prospero, 1982). There
are only a few studies to distinguish the proportions of
natural and anthropogenic components (Savoie et al.,
2002 is an example). New particle formation in the
atmosphere is inversely related to available aerosol
surface area (Clarke, 1993). So any sudden decrease in
aerosol concentration due to various removal processes
(especially precipitation) will result in the homogeneous
nucleation of the sulphur compounds. This leads to new
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particle formation in the marine boundary layer (MBL)
(Hoppel et al., 1994).
An increase in the marine DMS emission increases
the number density of sulphate aerosol over the
marine atmosphere and consequently the number
density of cloud droplets, which results in an increase
in cloud albedo (Charlson et al., 1987; Hegg, 1990;
Hegg et al., 1991; Covert et al., 1992, 1996). This
enhancement in cloud albedo will act as negative forcing
on global temperature; lower temperature in turn
results in reduced productivity and emission of marine
DMS. Charlson et al. (1987) estimated that a 40100%
increase in CCN concentration is sufficient to counter-
balance the temperature increase due to doubling of CO2concentration.
The organic, inorganic, mineral content and mass
concentration of the sub-micrometre aerosol were
measured in June and July 1997 on Tenerife in the
MABL and free troposphere (Putaud et al., 2000). Theyobserved that in the unperturbed MABL the aerosol
average composition was 37% non-sea salt sulphate,
21% sea salt and 20% organic carbon. In the
unperturbed free troposphere, organic carbon and
non-sea salt sulphate accounted for 43% and 32% of
the sub-micrometre aerosol mass respectively (Putaud
et al., 2000; Schmeling et al., 2000). Based on extensive
observations at the MABL simultaneously with the free
troposphere (FT), these studies have concluded that the
source for the free troposphere could be transport from
continents; in background conditions MABL aerosol is
formed by dilution of continental aerosol by FT air
modified by deposition and condensation of species of
oceanic origin. However, the outbreaks in the MABL
were due to transport of polluted air masses from
Europe.
The evolution of the aerosol characteristics in the
marine atmosphere was thoroughly studied during
Lagrangian experiments of ACE-2 (Johnson et al.,
2000). Observations during the first ACE-2 Lagrangian
experiment suggested that the important processes
controlling the sub-micrometre mode aerosol concentra-
tion, which dominated the total aerosol concentration,
included scavenging of interstitial aerosol by cloud
droplets, enhanced coagulation of Aitken mode andaccumulation mode aerosols due to increased sea salt
surface area and the dilution of MBL by FT air (Raes,
1995; Johnson et al., 2000). Observations during the
second ACE-2 Lagrangian experiment found evidence
of processing of aerosol particles by stratocumulus
cloud, in particular by aqueous phase reactions (Clarke,
1998; Osborne et al., 2000; Wood et al., 2000).
Measurements indicate that the concentration of
DMS is higher in summer than in winter and highest
over low-latitude oceans (Andreae, 1985; Bates et al.,
1987). These indicate that production of DMS increases
with an increase in ocean temperature, which depends
on the duration of sunlight received by the ocean
surface. The warmest, most saline and most intensely
illuminated regions of oceans have the highest rate of
DMS emissions to the atmosphere (Russell et al., 1994).
The largest DMS flux comes from the tropical and
equatorial oceans (Russell et al., 1994). The concentra-
tion of non-sea salt sulphates decreases from coastal
regions of the continent to the remote ocean areas
(Parungo et al., 1987; Fitzgerald, 1991).
3.3. Soil dust aerosols
Particles originating from the soil are usually mineral
aerosols and are produced by weathering of soil
(Jaenicke, 1980, 1993; Prospero et al., 1983, 2002;
dAlmeida, 1986; Zender et al., 2003; Ginoux et al.,
2004; Miller et al., 2004; Tegen et al., 2004). Ultra-fine
sand particles are formed by winds mostly in the arid
regions of the world (Pye, 1987; Schwartz et al., 1995;Prospero et al., 2002; Ginoux et al., 2004). The long-
range transport of continental derived particles by the
combined action of convection currents and general
circulation systems make these particles a significant
constituent even at locations far from their sources
(Delany et al., 1967; Prospero et al., 1970, 1981; Carlson
and Prospero, 1972; Prospero, 1979; Shaw, 1980;
Bergametti et al., 1989; Tegen and Fung, 1994; Arimoto
et al., 1995, 1997; Moorthy and Satheesh, 2000; Arimoto
et al., 2001; Zender et al., 2003; Ginoux et al., 2004). Soil
derived particles are among the largest aerosols with
radii ranging from below 0.1 mm to $100mm. Particles in
the size range r45 mm are present only in the source
regions but in general particles in the radius range
0.15mm are transported long distances ($5000km) into
the marine atmosphere (Arimoto et al., 2001; Prospero
et al., 2002; Gong et al., 2003; Maring et al., 2003; Reid
et al., 2003a,b). The measurements of aerosol size
distribution and analysis of chemical composition of
aerosols over Antarctica have found mineral particles
with radii greater than 2 mm of Australian origin (Shaw,
1980). The data from the TOMS satellite have been
extensively used to study the global distribution of dust
aerosols (Prospero et al., 2002). Global maps of TOMS
absorbing aerosol index shows an example of asignificant amount of dust aerosols over the Sahara
during the month of May (Prospero et al., 2002). When
the wind pattern is favourable these aerosols are
transported over the Atlantic Ocean and the Arabian
Sea to reach far ocean locations (thousands of kilo-
metres away from source).
There are a number of investigations available in the
literature regarding the transport of aerosols from
continents to ocean and vice versa (Eriksson, 1959,
1960; Toba, 1965a, b; Junge, 1972; Delany et al., 1973;
Prospero, 1979; dAlmeida, 1986; Bergametti et al.,
1989; Arimoto et al., 1995; Gong et al., 2003; Zender
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et al., 2003). Some of these authors found the existence
of Saharan dust even over the remote areas of the
Atlantic and Pacific Oceans (Carlson and Prospero,
1972; Junge, 1972; Prospero and Carlson, 1972; Pros-
pero, 1979; dAlmeida, 1986; Bergametti et al., 1989;
dAlmeida et al., 1991). Prospero et al. (1970) traced the
origin of a dust event at Barbados to West Africa with a
transport time of $5 days. The chemical analysis of
marine aerosol samples collected over the Atlantic
Ocean revealed an African source (Bergametti et al.,
1989). The major source of mineral dust in Africa is the
Sahara. Junge (1972) estimated that 60200 Tg Saharan
dust is generated over the Sahara and is transported
each year, whereas Duce et al. (1991) estimated that
$220 Tg mineral dust is transported to the North
Atlantic each year.
An example of dust transport over the Arabian Sea is
shown in Fig. 6a (using data from the moderate
resolution imaging spectro-radiometer (MODIS) on
board the TERRA satellite). The Arabian Sea region
has a unique weather pattern on account of the Indian
monsoon and the associated winds that reverse direction
seasonally. Chemical analysis of aerosols over the
tropical Indian Ocean have shown that more than six
months every year natural aerosols contribute more
than 50% to composite aerosol optical depth (Fig. 6b)
(Satheesh and Srinivasan, 2002; Satheesh et al., 2002).
They have demonstrated that radiative forcing due to
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49
71
5761
51
34
47
32
40 39
66
20
30
40
50
60
70
80
F M A M J J A S O N D
%ContributioninForcing
Natural Anthropogenic
30N
50E 55E 60E 65E 70E 75E 80E 85E 90E 95E 100E
0 0.1 0.2 0.3 0.4 0.5 0.6 0.7 0.8 0.9
(July 2003)Aerosol Optical Thickness
1
27N
24N
21N
18N
15N
12N
9N
6N
3N
E0
Fig. 6. (a) Aerosol optical depths over Arabian Sea demonstrating the transport of dust aerosols from Arabian Peninsula to Indian
region. (b) Contribution of natural aerosols to optical depth at Indian Ocean.
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natural aerosols in this region is about 1.5 times larger
compared to that due to anthropogenic aerosols. Most
of the natural aerosol forcing was contributed by dust
(from the Arabian Peninsula) and locally generated sea
salt. These observations are inconsistent with that
reported by Li and Ramanathan (2002).
For absorbing aerosols like dust, radiative forcing at
the surface differs substantially from the value at the
TOA and the climate response depends not only upon
the TOA forcing, but its difference with respect to the
surface value, which represents radiative heating
within the atmosphere (Miller et al., 2004). Surface
forcing alters evaporation and the hydrologic cycle.
Studies by Miller et al. (2004) have shown that while
global evaporation and precipitation are reduced
in response to surface radiative forcing by dust,
precipitation increases locally over desert regions, so
that dust emission can act as a negative feedback to
desertification.Dust aerosols are significant contributors to radiative
warming below 500 mb due to short-wave absorption
but they have less effect on long-wave radiation
(Mohalfi et al., 1998; Alpert et al., 1998; Miller and
Tegen, 1999 and Fig. 7). Typically, dust approximately
doubles the short-wave radiation absorption under
clear-sky conditions (Tegen and Miller, 1998). Tegen
and Fung (1994) has shown that dust from disturbed soil
causes a net cooling at the surface, accompanied by an
increase in atmospheric heating. Such radiative effects
are found to be most pronounced over the desert regions
(Mohalfi et al., 1998). There have been several investiga-
tions to understand the characteristics of the dust layer
and the radiative heat balance. There are only very few
studies on the impact of dust on synoptic-scale systems.
The reduction of solar radiation reaching the Earths
surface as a result of scattering and absorption by dust
aerosols reduces the sensible heat flux. This is balanced
by the radiative heating of dust aerosols at low levels.
The dust aerosols over the Arabian Sea warm the levels
between 800 and 600 hPa ($0.2Kday1) and cool the
lower levels during daytime (Alpert et al., 1998; Mohalfi
et al., 1998). Thus the presence of dust transported over
oceans intensifies a low-level inversion, which in turn
affects the stability of the atmosphere (Miller and Tegen,
1999; Mohalfi et al., 1998).
Both land and sea are heated during daytime by
radiation from the Sun. But since solar radiation onlypenetrates a few centimetres of soil so that only top layer
heats up. The air above heats up much more rapidly
because of the low heat capacity of air. On the other hand,
the sea warms up much more slowly because of the large
heat capacity as well as longer penetration of solar
radiation. Warm air rises over land causing a low-pressure
region compared to the ocean. To compensate for this, air
flows from sea to landthe well-known sea breeze. When
the winds are strong enough, the land areas (especially
with low vegetation cover) produce soil dust aerosols. The
presence of this dust reduces the surface-reaching solar
radiation due to scattering and absorption, and heats the
lower atmosphere due to absorption. This cooling from
below and heating aloft creates low-level inversion. This
reduces the intensity of convection currents, and thus
increases the atmospheric stability. The reduction of solar
radiation at the surface reduces the surface heating which
in turn decreases the landsea temperature contrast and
consequently the intensity of the sea breeze. Thus,
depending on the concentration of the dust layer, the
impact can be different. There can be changes in sea-breeze
onset time also.
Since stable conditions resist upward movement, we
might conclude that clouds would not form when stable
conditions prevail in the atmosphere. Since the surfaceair is cooler and heavier than the air aloft, little vertical
mixing occurs between layers. Since air pollutants are
added from below, temperature inversion confines them
to the lowermost layers where they continuously build in
concentration (Mohalfi et al., 1998). The fact that
atmosphere is either stable or not, determines whether
clouds develop or not. The accumulation of aerosols at
lower levels would increase the lower atmosphere
heating further, which in turn would increase the
stability (positive feedback).
Dust aerosols absorb sunlight to a greater extent than
industrial sulphate and sea salt aerosols (Tegen et al.,
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Fig. 7. The radiative impact of dust aerosol in short wave and
long wave regions. Dust aerosols larger in size and have
absorbing property in infrared and hence unlike other aerosol
species, dust aerosol influence infrared as well. The symbols SW
and LW represent short wave and long wave radiation and the
subscripts TD, TU, BD, BU represents top of the atmosphere
down-welling, top of the atmosphere upwelling, bottom of the
atmosphere (surface) down-welling, and surface upwelling,
respectively.
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1997; Miller and Tegen, 1999; Haywood and Boucher,
2000). These authors also suggest that dust optical
properties (to which the top of the atmosphere forcing is
sensitive) in global models should be allowed to vary
with the mineral composition of the source region in a
computation of the climate response. More extensive
measurements of the dust optical properties, along with
the vertical distribution of the dust layer, are needed to
reduce the uncertainty of the climate response to dust
aerosols.
Liao and Seinfeld (1998) examined radiative forcing
by mineral dust aerosols in short-wave and long-wave
regions using a one-dimensional column radiation
model. They estimated clear sky TOA radiative forcing
as 2 W m2 at low surface reflection ($0.1) and
+ 2 W m2 at high surface reflection ($0.5). Under
cloudy skies these values are in the range of +2 to
+ 3 W m2. They also observed that unlike scattering
aerosols such as sea salt, dust radiative forcing dependson the surface reflection, the altitude at which the dust
layer is located and the relative altitude from the cloud
layer. Clear sky TOA long-wave radiative forcing was in
the range of +0.21.0W m2 and corresponding values
for cloudy skies were 0.0 and +0.6 W m2. These results
are consistent with Tegen and Lacis (1996) and Tegen
et al. (1997).
Dust can serve as a reaction surface for reactive gas
species in the atmosphere (Dentener et al., 1996;
Huebert et al., 2003; Carmichael et al., 2003; Seinfeld
et al., 2004). Mineral dust is believed to play an
important role in marine biological processes (Maher
and Dennis, 2001; Prospero et al., 2002). Trace metals
on dust are essential to some marine biological
processes; for example, dust is a source of iron, which
acts as a nutrient for phytoplankton (Falkowski et al.,
1998; Fung et al., 2000; Maher and Dennis, 2001;
Prospero et al., 2002; Huebert et al., 2003; Carmichael
et al., 2003; Seinfeld et al., 2004).
3.3.1. Dust transport models
There is a substantial transport of mineral aerosol
from Asia to wide areas of the North Pacific with an
estimated total annual input in the range of 610
million tons year1. This atmospherically transporteddust is a significant source of sedimentary material for
the North Pacific. Global dust distributions are usually
calculated with transport models. Measurements of dust
at various locations alone cannot provide information
on its transport and consequent impact over other
regions. Mathematical models provide the necessary
framework for the integration of our understanding of
various atmospheric processes and to study their
interactions (Luo et al., 2003; Gong et al., 2003; Zender
et al., 2003; Ginoux et al., 2004; Tegen et al., 2004).
Measurements and models together provide a powerful
tool to study the dust aerosol transport. Many global
models do not accurately simulate regional distribution
of dust due to their low grid resolution and inaccuracy
of dust source function. To accurately predict the impact
of dust aerosols on climate the spatial and temporal
distribution of dust is essential. The dust emission is
calculated depending on soil moisture, surface wind
speed and soil surface conditions. The major sink is
gravity settling. The model simulations have shown that
the contribution of dust to aerosol optical depth is
927% for 201S201N, in general, 4066% in the Sahel
region and 3054% in East Asia (Tegen, 1994). Over the
Indian Ocean dust contributes 15% to total aerosol
optical depth during winter (Satheesh et al., 1999).
However, regional characteristics of soil dust produc-
tion, transport and removal processes are poorly
understood.
Recent studies have demonstrated that a fraction of
the atmospheric dust load originates from anthropo-
genically disturbed soils (Tegen et al., 2004). Bycalibrating a dust source model with emission indices
derived from dust storm observations, Tegen et al.
(2004) estimated the contribution to the atmospheric
dust load from agricultural areas to be o10% of the
global dust load. Comparisons between a 22-year
simulation of mineral aerosols with satellite and in situ
observations suggest that the model can predict atmo-
spheric mineral aerosol distributions, with some dis-
crepancies (Luo et al., 2003). In addition, there were
differences between the model results and previously
published results (e.g., Ginoux et al., 2001). The
sensitivity analysis showed that differences between
simulated dusts near Australia are likely due to
differences in both source parameterisation and surface
winds (Luo et al., 2003).
Zender et al. (2003) described a model for predicting
the size-resolved distribution of atmospheric dust for
climate and chemistry-related studies. The dust distribu-
tion from 1990 to 1999 is simulated with our mineral
aerosol entrainment and deposition model embedded in
a chemical transport model (Zender et al., 2003).
Without invoking anthropogenic mechanisms the model
captures the seasonal migration of the transatlantic
African dust plume, and it captures the spring maximum
in Asian dust outflow and concentration over thePacific. Zender et al. (2003) estimated the 1990s global
annual mean and variability of dust (diameter,
Do10mm) to be the following: emissions,
14907160 Tg yr1; burden, 1772 Tg; and optical depth
at 0.63 mm, 0.03070.004. These values for emission,
burden, and optical depth are significantly lower than
some recent estimates. The model underestimates trans-
port and deposition of East Asian and Australian dust
to some regions of the Pacific Ocean.
Gong et al. (2003), using a size-segregated soil dust
emission and transport model, Northern aerosol regio-
nal climate model (NARCM), simulated the production
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and transport of Asian soil dust during the aerosol
characterization experiment-Asia (ACE-Asia) period
from March to May 2001. The model was driven by
the NCEP reanalysed meteorology and has all the
atmospheric aerosol physical processes of soil dust:
production, transport, growth, coagulation, and dry and
wet deposition. Model simulations were compared with
ground-based measurements in East Asia and North
America and with satellite measurements for the same
period of time. The model captured most of the dust
mobilisation episodes during this period in China and
reasonably simulated the concentrations in source
regions and downwind areas from East China to western
North America. About 252.8 megatonnes of soil dust
below Do40mm was estimated to be emitted in the East
Asian deserts between 1 March and 31 May 2001.
Ginoux et al. (2004) simulated the global distribution
of aeolian dust from 1981 to 1996 with the global ozone
chemistry aerosol radiation and transport (GOCART)model. The simulated annual emission varies from a
minimum of 1950 Tg in 1996 to a maximum of 2400 Tg
in 1988. Of these emissions, 65% are from North Africa
and 25% from Asia. It was found that North America
received twice as much dust from other continents than
it emits per year. The inter-annual variability of dust
distribution was analysed over the North Atlantic and
Africa. It was found that in winter a large fraction of the
North Atlantic and Africa dust loading correlates with
the North Atlantic Oscillation (NAO) index. It is shown
that a controlling factor of such correlation can be
attributed to dust emission from the Sahel. However, the
long record of dust concentration measured at Barbados
indicates that there is no correlation with the NAO
index and surface concentration in winter. Longer
simulation should provide the information needed to
understand whether the effects of the NAO on dust
distribution are rather limited or whether Barbados is at
the edge of the affected region.
4. Radiative impact: natural versus anthropogenic
aerosols and GHGs
In this section, we compare the radiative forcing dueto various (natural and anthropogenic) aerosol species
as well as that due to GHGs.
4.1. Direct effect
Observations over the tropical Indian Ocean have
shown that TOA forcing due to sea salt aerosol is
1.3670.46Wm2 and that due to dust and soot are,
respectively 0.7270.3 and +0.6470.38Wm2. The
radiative forcing due to sulphate (natural and anthro-
pogenic) aerosol was 6.4Wm2. Haywood et al.
(1997), using a radiation code within a GCM, assessed
the direct radiative forcing by two major anthropogenic
aerosol components: anthropogenic sulphate and soot
aerosols from fossil fuel burning. They estimated that
under cloudy skies, radiative forcing due to anthropo-
genic sulphate is 0.6Wm2 for the northern hemi-
sphere and 0.15Wm2 for the southern hemisphere.
Similar results have been reported by Haywood and
Shine (1995), who report radiative forcing of
0.55Wm2 for the northern hemisphere and
0.13Wm2 for the southern hemisphere. For clear
skies, Haywood et al. (1997) reported a radiative forcing
of 0.59Wm2 for northern hemisphere and
0.14Wm2 for the southern hemisphere, which are
comparable with cloudy sky values. In the case of soot
aerosols, Haywood et al. (1997) estimated a radiative
forcing of +0.35 W m2 for the northern hemisphere
and +0.06Wm2 for the southern hemisphere under
cloudy skies. The corresponding values under clear skies
were +0.11 and +0.02W m2
.Haywood et al. (1999) have estimated clear sky
radiative forcing due to natural sulphate, natural dust
and sea salt as 0.93, 0.58, and 1.51Wm2 (for low
sea salt; 5.03Wm2 for high sea salt), respectively.
This means that radiative forcing due to natural aerosols
is 3.02Wm2 (for low sea salt; 6.54Wm2 for high
sea salt). They estimated the corresponding values for
anthropogenic sulphate, organic carbon, black carbon
and anthropogenic dust as 0.72, 1.02, +0.17,
0.54Wm2, respectively. Thus radiative forcing due
to anthropogenic aerosols is 2.11Wm2. These results
clearly show the significant role natural aerosols have in
determining the radiative forcing due to a composite
aerosol system.
Using an aerosol transport model coupled with a
GCM, Tekemura et al. (2002) estimated radiative
forcing due to various aerosol species. The global mean
radiative forcing due to black carbon under cloudy skies
was +0.36 W m2 and that due to anthropogenic
sulphate was 0.32Wm2. The corresponding values
for clear sky conditions were +0.21 and 0.72Wm2,
respectively. These values are slightly smaller than those
estimated by Penner et al. (1998) and Kiehl et al. (2000),
but comparable with those estimated by Boucher and
Anderson (1995) and Feichter et al. (1997). Sea salt anddust radiative forcing were +0.36 and 0.31Wm2
under cloudy skies and +0.26 and 0.59Wm2 under
clear sky conditions. The value of dust radiative forcing
is higher than the value of +0.14 W m2 reported by
Tegen et al. (1996).
The estimate of Tekemura et al. (2002) of radiative
forcing due to organic carbon, black carbon and
anthropogenic sulphate (total anthropogenic forcing
of 0.96Wm2) is comparable with that of Haywood
et al. (1999) when considering that anthropogenic dust
was not included in Tekemura et al. (2002). They did not
provide forcing due to natural sulphate. If we use the
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value of natural sulphate forcing from Haywood et al.
(1999), radiative forcing due to natural aerosols is
1.26Wm2.
An atmospheric general circulation model is coupled
to an atmospheric chemistry model to calculate the
radiative forcing by anthropogenic sulphate and carbo-
naceous aerosols (Penner et al., 1998). They estimated
that the direct forcing by anthropogenic sulphate
aerosols is in the range of 0.55 to 0.81Wm2. The
climate forcing associated with fossil fuel emissions of
carbonaceous aerosols is calculated to range from
+0.16 to +0.20W m2. The direct forcing of carbonac-
eous aerosols associated with biomass burning is
calculated to range from 0.23 to 0.16Wm2. Myhre
et al. (1998) estimated that the direct radiative forcing
due to sulphate and soot is 0.32 and +0.16Wm2,
respectively.
The above discussion shows that the radiative forcing
due to sea salt aerosols ranges from 0.5 to 6.0Wm2
while that of natural dust aerosols ranges from 2 to
+0.5Wm2. Now, we discuss the IPCC (2001) esti-
mates of the radiative forcing due to anthropogenic
aerosols. The global mean direct radiative forcing due to
anthropogenic sulphate aerosols reported by IPCC
ranges from 0.26 to 0.82Wm2 based on several
studies (Kiehl and Briegleb, 1993; Boucher and Ander-
son, 1995; Feichter et al., 1997; Graf et al., 1998;
Haywood et al., 1997; Hansen et al., 1998; Haywood
and Ramaswamy, 1998). The IPCC estimates of black
carbon (BC) aerosols from fossil fuel and biomass
burning is in the range +0.27 to +0.54 W m2, and the
corresponding estimate for organic carbon (OC) is in the
range 0.04 to 0.41Wm2 (Hansen et al., 1998;
Jacobson, 2001). It should be noted that uncertainties in
these estimates are large due to the limited number of
studies available.
Next, we come to radiative forcing due to GHGs.
Myhre et al. (1998) have performed calculations of the
radiative forcing due to changes in the concentrations of
the most important well-mixed GHGs since pre-indus-
trial time, and found that the radiative forcing due to all
the well-mixed GHGs is +2.25 W m2. IPCC reports
that radiative forcing due to major GHGs such as CO2,
CH4, N2O is +1.46, +0.48 and +0.15 W m2, respec-tively. The total radiative forcing due to well-mixed
GHGs is 2.43W m2. Thus negative forcing by naturally
occurring aerosols is quite significant when we consider
the fact that forcing caused by projected doubling of
CO2 is about +4 W m2 (Charlson et al., 1992; Winter
and Chylek, 1997).
A comparison of the radiative forcing due to various
aerosol species with that of GHGs is shown in Fig. 8a
(data obtained from the literature discussed in Sections 2
and 3 and summarised in Table 2). It can be seen that
sea salt aerosol forcing (and its variability) is quite large
compared to other species.
4.2. Indirect effect
Sea salt aerosols and natural sulphates are hygro-
scopic in nature and hence act as condensation nuclei for
the formation of clouds (Fitzgerald, 1991). Cloud albedohas a significant role in determining the global energy
balance (Chuang et al., 1997). An increased concentra-
tion of aerosols results in an enhanced concentration of
cloud droplets, which in turn increases the albedo of
clouds and this causes a decrease in the short-wave solar
radiation reaching the Earths surface (Clarke, 1998).
The increase in condensation nuclei (CN) also influences
the cloud lifetime. An increase in CN increases the cloud
droplet concentration and reduces the mean droplet size.
This increases the cloud lifetime and inhibits precipita-
tion. This also leads to an increase in fractional cloud
coverage and influences both short-wave and long-wave
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-6.5
-5.5
-4.5
-3.5
-2.5-1.5
-0.5
0.5
1.5
2.5
3.5
GHGs Sea-salt Dust BC OC Sulphate
(N)
Sulphate
(A)
TOADirectForcing(Wm
-2)
Global Average
Indian Ocean : Regional
-10
-9-8
-7-6
-5
-4-3
-2-1
0
Anthropogenic Sea-salt Direct Sea-salt Indirect
TOAForcing(W
m-2)
Anthropogenic Forcing = -5 2.5 W m-2
Sea-salt Direct Forcing = -2 1 W m-2
Sea-salt Indirect Effect = -7 4 W m-2
Fig. 8. (a) Comparison of greenhouse gas forcing with that of
aerosol forcing due to various species. (b) Natural vests
anthropogenic forcing over tropical Indian Ocean [The data
from the following sources: Kiehl and Briegleb, 1993; Boucher
and Anderson, 1995; Tegen and Lacis, 1996; Feichter et al.,
1997; Graf et al., 1998; Haywood et al., 1997; Tegen et al., 1997;
Moorthy et al., 1997; Winter and Chylek, 1997; Alpert et al.,1998; Mohalfi et al., 1998; Miller and Tegen, 1999; Haywood
and Ramaswamy, 1998; Penner et al., 1998; Haywood et al.,
1999; Satheesh and Ramanathan, 2000; Podgorny et al., 2000;
Jacobson et al., 2001; Ramanathan et al., 2001; Satheesh, 2002;
Tekemura et al., 2002; Soden et al., 2002; Satheesh and Lubin,
2003; Vinoj and Satheesh, 2004].
S.K. Satheesh, K. Krishna Moorthy / Atmospheric Environment 39 (2005) 20892110 2101
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radiation. Cloud albedo depends on the cloud droplet
number. For a given water vapour content, the average
cloud droplet size is larger for a lower number of
aerosols and is smaller for a higher number of aerosols
(Han et al., 1998). This is because the water vapour
availability per CN is more in the former case compared
to latter case. But the relation between aerosol number
and number of cloud droplets is not simple and depends
on a number of factors, including the aerosol chemical
composition, size distribution, supersaturation of air
and so on (Clarke, 1993; Ramanathan et al., 2001). Not
all aerosols are capable of acting as CN. To be able to
act as CN, the aerosol should be larger than a criticalsize ($1 mm) and should be hygroscopic (water-soluble)
(Hoppel et al., 1990, 1994). As the number of aerosols
increases, the supersaturation (S) reduces. The inverse
correlation is due to the fact that as more drops form,
the water supply available will be less and as a result Sis
reduced (Ramanathan et al., 2001).
Based on direct measurements of aerosols, cloud
droplet concentration and supersaturation over the
tropical Indian Ocean, Ramanathan et al. (2001) derived
empirical relations between aerosol number and various
parameters such as cloud drop number, cloud drop
effective radius, cloud optical depth and so on. Their
basic equation is of the form,
NCCN 0:12N1:25S=30:76, (3)
where NCCN is the number of aerosols which are
activated, N is the total number of particles, and S is
the supersaturation in percentage. The equation is valid
for values of N ranging from 300 to 2000 cm3 and
So0.3%. Here S is a function of N as the amount of
water vapour available per nuclei depends on the total N
for a given water vapour amount. They also found from
observed data that not all CCN becomes cloud droplets.
When the total aerosol number is low almost all CCN
becomes cloud droplets, whereas at high aerosolconcentrations, only about 80% of the CCN becomes
cloud droplets. The effective radius of cloud droplets
decreases from $8.0 to $5.5 when aerosol numbers
change from 300 to 2000cm3 (Ramanathan et al.,
2001). The number of cloud droplets increases from $75
to $300 cm3 and the corresponding cloud optical depth
increases from $3 to 14 for the same change in aerosol
number (Ramanathan et al., 2001).
Investigations have revealed that sea salt number
concentration over the ocean is a function of wind speed
(Lovett, 1978; Blanchard and Woodcock, 1980; ODowd
and Smith, 1993; Parameswaran et al., 1995; Moorthy
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Table 2
Comparison of direct radiative forcing (W m2) by various species
Location (regional/global) Species Radiative forcing (W m2) Reference
Global Sea salt 0.6 to 2.0 (low) Winter and Chylek (1997)
1.5 to 4.0 (high)
Global Sea salt 1.51 (low) Haywood et al. (1999)
5.03 (high)
Deserts Dust 2 to +2 Liao and Seinfeld (1998)
Indian Ocean Sea salt 1.3670.5 Podgorny et al. (2000)
Indian Ocean Sea salt 1.5 to 6.0 Satheesh and Lubin (2003)
Indian Ocean Dust 0.7270.3 Podgorny et al. (2000)
Indian Ocean Soot (BC) +0.6470.4 Podgorny et al. (2000)
Indian Ocean Sulphate (natural and anthropogenic) 6.470.5 Podgorny et al. (2000)
Northern Hemisphere Sulphate (anthropogenic) 0.55 to 0.6 Haywood and Shine (1995)
Haywood et al. (1997)
Southern Hemisphere Sulphate (anthropogenic) 0.13 to 0.15 Haywood and Shine (1995)
Haywood et al. (1997)
Northern Hemisphere Soot (BC) +0.11 Haywood et al. (1997)
Southern Hemisphere Soot (BC) +0.02 Haywood et al. (1997)Global Sulphate (anthropogenic) 0.72 Haywood et al. (1999)
Global Soot (BC) +0.17 Haywood et al. (1999)
Global Sulphate (natural) 0.58 Haywood et al. (1999)
Global Dust 0.93 Haywood et al. (1999)
Global Sulphate (anthropogenic) 0.72 Tekemura et al. (2002)
Global Soot (BC) +0.21 Tekemura et al. (2002)
Global Dust +0.14 Tegen et al. (1996)
Global Sulphate (anthropogenic) 0.26 to 0.82 IPCC (2001)
Global BC (FFB) +0.27 IPCC (2001)
Global BC (BB) +0.57 IPCC (2001)
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et al., 1997). ODowd and Smith (1993) reported that sea
salt number increased by 100cm3 when wind speed
increased from 3 to 15 m s1. This observation, when
combined with observations of Ramanathan et al.
(2001), shows that a change in wind speed from 3 to
1 5 m s1 can change the cloud droplet number by
$30cm3 and increase the cloud optical depth by $3.
The estimates of sea salt direct and indirect effects over
the Indian Ocean were 271 and 774 W m2,
respectively (Vinoj and Satheesh, 2004). This is quite
large compared to anthropogenic aerosol forcing
reported over this region (572.5Wm2) (Rama-
nathan et al., 2001). Thus, clearly the direct and indirect
effects of sea salt aerosols have a significant role in
offsetting the positive forcing by absorbing aerosols
and GHGs.
The direct and indirect forcing due to sea salt aerosols
compared with anthropogenic forcing over the Indian
Ocean is shown in Fig. 8b. The magnitude of indirectradiative forcing (and uncertainty) due to sea salt
aerosols is several-fold more than the direct radiative
forcing of sea salt aerosols. The large magnitude and
variability in both direct and indirect forcing due to sea
salt aerosols emphasises the importance of natural
aerosols.
Soil dust is not hygroscopic and as such does not
participate as CCN. There are two extremes of insoluble
nuclei: nuclei which are activated (wetted) easily, and
nuclei which are not easily activated. Nuclei which are
easily activated rapidly, get coated with liquid and
subsequently behave like droplets and further grow in
size by condensation (Levin et al., 1996; Wurzler et al.,
2000; IPCC, 2001). The droplet growth thereafter can be
predicted by using Kelvins equation. In cases where
nuclei surfaces are not wettable, condensation proceeds
with much more difficulty. The surfaces of the nuclei try
to make the condensing liquid into small spheres. When
the entire surfaces are covered with these small spheres,
liquid coatings can form. Hereafter the nuclei behave
like normal droplets and grow in size by condensation of
vapour. Soil dust is often internally mixed with other
species and thus can be hygroscopic (Prospero et al.,
2002). Levin et al. (1996) observed that desert dust was
coated with sulphate, which probably originated fromin-cloud scavenging of interstitial dust particles followed
by evaporation of the cloud droplets. The presence of
soluble materials over dust makes them into large and
effective CCN, which may affect cloud microphysics
(Levin et al., 1996; IPCC, 2001). The role of insoluble
nuclei in condensation is still a question to be answered
(Levin et al., 1996; Wurzler et al., 2000).
The IPCC estimates of the indirect radiative effect due
to anthropogenic sulphate ranges from 0.3 to
1.8Wm2 based on various studies (Chuang et al.,
1997; Boucher and Lohman, 1995; Jones and Slingo,
1996, 1997). Chuang et al. (2002) obtained an indirect
radiative forcing due to black carbon and organic
carbon aerosols of 1.51Wm2. Kaufman and Naka-
jima (1993) have estimated the indirect radiative forcing
by smoke to be 2 W m2 using satellite data over
Brazil.
5. Summary and conclusions
Aerosols are of natural or anthropogenic origin.
Natural aerosols account for $70% of the global
aerosol loading and of this the main contributors are
sea salt, dust and natural sulphates. Nevertheless, the
abundance of these shows significant variability from
region to region and season to season. Recent investiga-
tions have shown that a proportion of dust is due to
anthropogenic activities. Similarly, a proportion of
anthropogenic soot originates from natural forest fires.
It is difficult to separate the anthropogenic componentsof dust from natural, or natural components of soot
from anthropogenic. Besides, away from the source
regions, both natural and anthropogenic components
mix together and on a global scale it is almost impossible
to exactly apportion the natural and anthropogenic
shares of the total aerosol. Nevertheless, several
investigations and coordinated field campaigns have
been carried out to assess the impact of anthropogenic
aerosols on climate (particularly because they are
amenable to mitigation). The ACE-2, Tropospheric
Aerosol Radiative Forcing Experiment (TARFOX),
Indian Ocean Experiment (INDOEX) are examples.
The ACE-1 and ACE-Asia, however, have provided
valuable information on natural aerosols. Even so, there
are still far fewer studies on natural aerosols compared
with anthropogenic aerosols, despite their importance.
To accurately predict the impact of dust aerosols on
climate, the spatial and temporal distribution of dust is
essential. However, regional characteristics of soil dust
production, transport and removal processes are poorly
understood. Many global models do not accurately
simulate regional distribution of dust due to their low
grid resolution and inaccuracy of dust source function.
To accurately predict the impact of dust aerosols on
climate the spatial and temporal distribution of dust isessential. More extensive measurements of the dust
optical properties, along with the vertical distribution of
the dust layer, are needed to reduce the uncertainty of
the climate response to dust aerosols. Similarly, there are
very few data on sea salt aerosols where wind speeds are
high. In such conditions accurate measurements are
extremely difficult. Thus the data on the global
distributions of two major natural aerosol types (sea
salt and mineral dust) are not adequate.
Several experiments and simulations have attempted
to quantify the radiative impacts of natural aerosols,
particularly sea salt, dust and oceanic sulphate, yet large
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uncertainties persist in these estimates especially due to
the following:
Fewer data over oceans, especially on sea saltaerosols, and limitations of available observations
over oceans. At high wind speeds, it is difficult tomake accurate measurements over oceans.
Inadequate understanding of optical/radiative prop-erties of dust aerosols, their large regional differences
(depending on the soil type at the source region), the
transport processes and sinks.
Thus, more data using well-focussed experiments are
needed to reduce the uncertainties in the characteristics
of these two major natural aerosol species from a global
perspective.
Notwithstanding the above, estimates have shown
that even at low wind speeds, radiative forcing due to seasalt aerosols can be in the range from 0.5 to 2 W m2
and at higher wind speeds this can be as high as in the
range 1 to 6 W m2. This negative forcing (cooling)
by naturally occurring sea salt aerosols is quite
significant when we consider the fact that the forcing
caused by the increase in CO2 since the advent of the
industrial era is about +1.46 W m2, and forcing caused
by projected doubling of CO2 is about +4W m2.
Similarly detailed estimates of the dust radiative forcing
shows values in the range of 2 to +0.5 W m2. It may
be noted that due to the poor data on regional
characteristics of soil dust source function, we do not
even know whether dust radiative forcing is positive ornegative. Most of the recent investigations, however,
indicate that dust radiative forcing is negative.
Thus, natural aerosols contribute quite significantly to
global radiative forcing. This contribution has large
seasonal and spatial variability and is comparable to or
even larger than anthropogenic forcing. Though no
steps can be taken to reduce these effects (unlike
anthropogenic effects, which are amenable to mitiga-
tion), a clear understanding is needed to appreciate the
climate impact of aerosols on the one hand and
the extent of perturbation caused by human activities
on the other hand.Thus to assess the climate impact of aerosols, while
separating out the human factor, we need to address the
following issues.
A. To what extent can wind-generated sea salt aerosols
offset the atmospheric heating due to absorbing
aerosols such as soot transported over oceans?
B. At high wind speeds, newly produced sea salt
droplets may coat over pre-existing absorbing soot
aerosols, thus significantly altering their absorbing
efficiency. What is the consequent impact on
radiative forcing? This could significantly alter the
aerosol properties not only over oceans but also over
a significant part of the continents along the vast
coastal areas of the globe.
C. The reduction in solar radiation at the surface
simultaneous with lower atmospheric heating by
dust aerosols could intensify a low-level inversion
and reduce the sensible heat flux. How does this
impact the formation of clouds?
D. The dust containing iron transported over the ocean
serves as nutrients to marine phytoplankton. The
consequent enhancement in DMS emission (due to
iron fertilisation) will increase the natural sulphate
aerosols over the ocean, which may have an
influence on cloud droplet concentration, cloud
albedo and hence alter the radiation balance as
much as, or at times even more than, the changes
brought about by anthropogenic sulphates over
oceans.
E. Dust optical properties vary from region to region.Recent investigations have shown that the dust
absorption is lower than that assumed in global
models. Regional distribution of dust source func-
tion is poorly understood due to lack of an adequate
database.
F. Prior to 2001, international panels such as the
Intergovernmental Panel on Climate Change (IPCC)
focussed mainly on the anthropogenic aerosol
components. In IPCC (2001), great effort was made
to assess the impact of natural aerosols. It is,
however, true that we lack sufficient information
on natural aerosols over large areas of the world,
especially the oceans.
G. The presence of natural aerosols influences the
anthropogenic aerosol forcing (either directly or
indirectly). I t is difficult to separate the natural and
anthropogenic aerosol contributions on radiative
forcing when they are in a mixed state.
Hence there is an urgent need to focus attention on
radiative effects of natural aerosols, especially in the
tropics where data on aerosols are sparse.
Acknowledgements
The authors thank ISRO for supporting this study
through the ISRO-Geosphere Biosphere Programme.
We thank Prof. J. Srinivasan, CAOS, Indian Institute of
Science, for valuable suggestions.
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