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Water Isotopes (δD and δ18O in the Hydrosphere) Reading:

White’s Lecture 20 Alley and Cuffy chapter in the Valley and Cole Book (pdf posted on the class web site)

Other resources: Ian Clark and Peter Fritz, Environmental Isotopes in Hydrology, Lewis, 1997. Web site: http://www.sahra.arizona.edu/programs/isotopes/oxygen.html#4 Some information can be found at the International Atomic Energy Agency web site: http://www-naweb.iaea.org/napc/ih/index.html http://www-naweb.iaea.org/napc/ih/documents/userupdate/description/1stpage.html

Motivation: Fractionation of oxygen and hydrogen isotopes occurs in the atmosphere and hydrosphere and gives us some very powerful tools for tracking climate change and hydrologic phenomena.

Guide Questions: What are the standards used for H and O isotope measurements? Is isotopic equilibrium approached during precipitation of water from H2O vapor? Why is precipitation a Rayleigh process? Is isotopic equilibrium approached during evaporation of water to form H2O vapor? Roughly how large are the isotopic fractionations for the water - H2O vapor equilibrium at 0˚C and 100˚C, for oxygen and hydrogen? Explain why the isotopic composition of precipitation on earth corresponds somewhat to a Rayleigh distillation model? How is the actual system on earth NOT like a simple Rayleigh distillation process. What is the rough relationship between δ18O in average precipitation and mean annual temperature for the entire earth? Why is this relationship strong, when δ18O actually depends on the amount of rainout? About how much δ18O change occurs, on average, per degree Celsius change in T? How much of the line plotted by Dansgaard (1964) is applicable to mid latitudes? How much scatter is there away from the best fit curve, in δ18O values? What is the global meteoric water line and what is its slope? Why do some bodies of water plot to the right of the GMWL? What process is indicated by this? What happens to the isotopic composition of groundwater during extensive water-rock interaction? Why? What patterns appear in the O and H isotopes of ice cores from Antarctica and Greenland? How do they compare with other measures of climate? Are the O and H isotopes in ice cores perfect paleothermometers? What are the potential problems? Are δD and δ18O in precipitation different for different seasons? Why? In a given area are δD and δ18O in precipitation different for different altitudes? Why? Why does a record of δD and δ18O in precipitation on, or perhaps downwind from, a mountain range tell us about paleoaltitude? How is δD and δ18O in groundwater that was recharged to an aquifer 20,000 years ago different from that of today (say, for Illinois). Is it possible for drinking water in a city or town to have different δD and δ18O compared to the local precipitation? Why?

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Standards used for oxygen and hydrogen isotopes Ocean water is used as the standard for all H isotope analyses and for most O isotope analyses. Actually, there is slight variation in ocean water, so we have something called: - Standard Mean Ocean Water (SMOW) as a standard. V-SMOW is an equivalent standard

developed by the IAEA in Vienna. The original SMOW not available any more. - PDB (now VPDB) is a carbonate material used for oxygen isotopes in carbonate rocks.

δ18O(VSMOW) = δ18O(VPDB) + 30.92 + δ18O(VPDB)* 30.92/1000

Fractionation of O and H: equilibrium and kinetic - Precipitation (rain or snow) is an equilib. process.: During precipitation of water from H2O vapor,

the intimate contact between water and vapor, and the fact that humidity is very close to saturation (100% relative humidity) (see appendix for an exception)

- Precipitation is a Rayleigh process: Precip. is removed from the cloud and thus back-reaction ceases and we have a Rayleigh process.

- Evaporation is kinetic (it may get close to equilibrium if humidity is high and thus back reaction is almost as great as forward reaction- but of course net evaporation is slow at high humidity)

- Sizes of equilibrium fractionations- 1000lnα values given as a function of T in a figure in White’s chapter. Actually that figure is confusing, so here’s a better one:

At 0˚C: 1000lnα ( ≈ δliquid - δvapor) = 11‰ for oxygen isotopes = 99‰ for hydrogen isotopes At 100˚C: = 3.2‰ for oxygen isotopes = 29‰ for hydrogen isotopes (From Faure, 1986)

A little more on the physics of water-vapor equilibrium In his lectures, White uses a model focusing on the vapor pressures of H2O’s bearing lighter vs. heavier isotopes. This gives us a reason to believe 1000‰lnα can be fit to the function: 1000‰lnα = A-B/T or 1000‰lnα =A-B/T2. I will not repeat his derivation, as you have already gotten the general idea of how isotopic fractionations are modeled.

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Application of Rayleigh fractionation and transport models to H2O vapor in the Earthʼs atmosphere (and precipitation). A: The classic model (has some flaws): 1) Most water vapor on earth originates in the tropics 2) It is then transported, via winds and mixing, toward the poles 3) The warm air in the tropics is moisture-laden 4) As a packet of air cools, it drops moisture as rain or snow 5) The transport of moisture poleward thus involves a distillation process

- remaining vapor becomes progressively depleted in heavier isotopes 6) The fraction of the original moisture remaining should be a function of temperature

- and thus, the isotopic depletion should be a function of temperature 7) Using a Rayleigh fractionation model, we can then link temperature to δ18O

NOTE: 1000lnα increases as T decreases, so this Rayleigh model is more complex than the one we have done for constant α. See Alley and Cuffey, Fig. 3 for the constant α case. Note that in White’s notes, he presents a model for fractionation of water vapor for a simplified case as an example only. Don’t use that model for the real world.

B: A more realistic picture involves the following:

1) “Rain-out” effect as given in the classic model above 2) Mixing of different air masses, with variable amounts of moisture coming from different

places 3) Limited exchange of H20 between atmosphere and oceans after the air leaves the tropics 4) Limited exchange of H20 between atmosphere and soils, lakes, and rivers 5) Orographic effects, e.g., Nevada, USA has isotopically light rain because of strong rainout

over the Sierra Nevada mountains of California. 6) Evaporative/Recycling effects: Lake/river water subjected to evaporation becomes more

depleted in heavy isotopes than the parent rain. Also operates in Nevada, possibly. 7) For interpretation of ancient conditions: Changes in Moisture source T and humidity. 8) For interpretation of ancient conditions: Changes in δ18O of oceans (about 1 per mil heavier

during last glacial period). - If you want to see, in more detail, realistic models that help us understand some of these

issues, see: - Hendricks et al. (2000) [Global Biogeochem. Cycles, Vol. 14, pp. 851-861] and - Jouzel et al., [JGR-D Vol. 92, 14739-14760] Actual δD and δ18O on earth (annual

average) as a function of mean annual air temperature.

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Relationship between mean annual T and δ18O (or δD)

To the left is the Dansgaard (1964) curve- Measurements of δ18O plotted versus mean annual temperature.

Also see Fig. 4a from Alley and Cuffey. Their plot is particularly useful, as it has different symbols for different latitude ranges. NOTE that most of the line is defined by data from Greenland and Antarctica. 1) Interesting that the data are close to linear. Theory gives some curvature. Actually, the plot in Alley and Cuffey looks somewhat curved. 2) However, there are some locations that fall far off the line. Especially above the line. This is caused by local, low-altitude moisture derived from the ocean (e.g., consider an island in the arctic, surrounded by ocean- where does the moisture come from?)

Other questions: Shouldn’t δ18O depend on the temp. at the time and place of precip, but NOT necessarily on the mean annual temp? Clearly, these two are closely related, so this is often a minor error. But in some areas most of the precipitation is in one season only and thus the mean annual temp. is not appropriate for the plot. A side note on how the earth works: Water vapor is MUCH MORE IMPORTANT in determining climate and weather than you might think. The strong relationship between temperature and water isotopes reflects an underlying connection between heat transport and water vapor on earth:

1. The latent heat in water vapor is the dominant mechanism by which energy gets transferred from the ocean/land surface to the atmosphere.

2. The latent heat in water vapor is the dominant mechanism by which energy gets transferred from low latitude to high latitude.

3. When precipitation occurs, the latent heat in water vapor is released to the air. a. Example: Air descending into Nevada from the Sierra Nevada Mtns of California

ends up warmer than it was prior to going over the mountains b. Example: A cold polar air mass collides with a moist air mass in Illinois. The

resulting storm drops snow and the cold air mass is greatly warmed. In short: Water vapor generation/movement/condensation is the most important heat transfer mechanism on earth and it controls climate to some extent.

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A little more on the physics of evaporation White’s lectures say very little about isotopic fractionation during evaporation of H2O. Evaporation is a kinetic process and the fractionation factors for O and H isotopes are different from the equilibrium ones. A good model is described in the Clark and Fritz book. They say isotopic fractionation during evaporation is driven by:

1) Equilibrium fractionation between the water surface and a thin layer of air (rheologically, the boundary layer) that is very close to 100% humidity, and

2) Kinetic effect: Faster diffusion of lighter isotopologues (e.g., HDO and H2O are isotopologues), relative to heavier ones out of this zone into the lower-humidity air above. Example: The diffusion of HDO (mass = 19) is slower than that of H2O (mass =18).

This approach has the advantage of explaining why humidity affects the fractionation factors (see below). Reference: Gonfiantini, 1986. R. Gonfiantini , Environmental isotopes in lake studies. In: P. Fritz and J.C. Fontes, Editors, Handbook of Environmental Isotope Geochemistry vol. 3, Elsevier, New York (1986), pp. 113–168.

Relationship between δD and δ18O on earth: The “meteoric water line” and how to detect evaporation. NOTE: White has nothing in his lectures on this topic. One good resource is: http://www.sahra.arizona.edu/programs/isotopes/oxygen.html

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1) The solid line gives the approximate compositions of rain/snow on earth. This is known as the Global Meteoric Water Line (GMWL). δD =(7.96) δ18O + (8.86) (Ronzanski et al., 1993) Rain at low latitude plots along the line to the left of the seawater point. Rain-out drives the remaining water vapor along the line to the lower left. Higher latitudes plot to the lower left. The GMWL is the effect of Rayleigh-type fractionation driven by equilibrium fractionation during precipitation.

2) Evaporation trends. The dotted line gives the trajectory followed (approximately) by a mass

of water (e.g., a lake) as a large fraction of it is evaporated away. The kinetic fractionation occurring during evaporation is different from equilibrium fractionation occurring during precipitation, so the slope (relative effects on hydrogen versus oxygen isotopes) is different. - The exact slope depends on the humidity (see plot below). Low humidity leads to a slope

very different from that of the GMWL. High humidity- slope more similar to GMWL.

Plot of δD vs. δ18O: Close linear relationship between δ18O and δD, as expected, because the relative sizes of the O and D fractionations should be constant (for precipitation; see below for evaporation). The diamond gives the composition of seawater. Tropical rains plot just to the left of the seawater value. Precipitation samples from cooler places plot along the GMWL; colder places are farther to the lower left. The horizontal arrow gives the effect of extensive water-rock interaction (e.g., hot spring waters).

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- - Evaporation from soils tend to be more like the high humidity case - Transpiration: Water uptake by plants does not involve much of an isotopic fractionation

and therfore the overall transpiration flux must NOT be highly fractionated isotopically

So... compositions falling to the right side of the line suggest that the water being measured have previously lost water vapor via evaporation. Example: Hydrologists can use this as a means of detecting lake water that has moved into the subsurface. 3) Deuterium Excess. Departure from the GMWL can be quantified by a parameter know as the “deuterium excess” = d or DE. This is defined as the y-intercept of the line with a slope of 8.0, drawn through a sample’s composition. The equation is:

d = δD – 8δ18O The deuterium excess can be used to: - Infer humidity at the moisture source. If most of the moisture in an air mass was transferred to

the atmosphere under low-humidity conditions, a larger d results. - Distinguish different moisture sources (different air masses coming into a region) that produced

recharge for old groundwater Joel Gat (Wiezmann Institute, Israel) has done much research on the meteoric water lines and the effects of evaporation. See also: Craig, H. and Gordon, L.I., 1965. Deuterium and oxygen-18 variations in the ocean and marine atmosphere. In: E. Tongiorgi (Editor), Stable Isotopes in Oceanographic Studies and Paleo-Temperatures. Here is a recent article on D excess applied to an ice core record: J. Jouzel, M. Stievenard, S.J. Johnsen, A. Landais, V. Masson-Delmotte, A. Sveinbjornsdottir, F. Vimeux, U. von Grafenstein, J.W.C. White, The GRIP deuterium-excess record, Quaternary Science Reviews, Volume 26, Issues 1-2, January 2007, Pages 1-17, ISSN 0277-3791, DOI: 10.1016/j.quascirev.2006.07.015. (http://www.sciencedirect.com/science/article/B6VBC-4MBCBPX-1/2/8c7e57f2e4d99e7992b70b831f514588)

4) Local Meteoric Water lines. Some places have precipitation that plots well away from the GMWL. In fact, every place has its own Local Meteoric Water Line (LMWL); some of these are very close to the GMWL, but many, especially, in arid regions, are not. For example, in Nevada and Israel, the climate is very dry, most of the water vapor in the air has been recycled through one or more cycles of precipitation then evporation, and evaporation effects drive the water vapor to unusual compositions. So we define a local meteoric water line and use that for local studies to identify meteoric waters- evaporation then shifts remaining waters in lakes and rivers away from the LMWL. 5) Effects of high-temperature chemical reaction between water and rock. The horizontal arrow in my graph above shows the effects of extensive water-rock interaction at high temperatures. This happens because the water exchanges oxygens with the host rock, which has relatively heavy oxygen isotope ratios. Hydrogen is affected little by this process, because there’s so much more hydrogen in water than in rock.

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δ1 8O and δD data in Ice cores: A record of climate change. Ice cores from Antarctica: Snow δ18O was 8 to 10 ‰ lighter during last glacial period- why? 1) Is it the T effect on the δprecip. − δvapor during precip.? NO. This site was likely a few degrees colder, so, if the vapor were the same then as now, the snow would be isotopically heavier during glacial time- REVERSE of observed 2) Rainout effect: This is the dominant effect. The vapor was lighter during the cold period, and this must be caused by greater rainout occurring prior to the air mass arrival at the site. Dansgaard model predicts that isotope ratio should be a strict function of temp (because the amount of rainout is). So one interpretation is that the δ18O shift tells us it was 13˚C colder and no other changes occurred. - Perhaps there was more rainout without a T change? Is this plausible? During colder periods,

we expect less vigorous circulation of the atmosphere, and possibly less efficient transport of moisture toward poles. But this requires either: - 1) Colder temp needed to get more rainout – T change required - 2) Different ”starting point” for the rainout- e.g., maybe rainout started at lower latitude?

Maybe conditions at the moisture source were different? . - 3) Different amount of mixing between air masses as poleward transport occurs?

3) Preservation effects? Maybe summer snow is ablated away more during warmer times- this causes a loss of some of the heaviest snow of the year. So colder times would have less of this and would be isotopically heavier- wrong direction, doesn’t explain observed shift. Summary: The temperature was almost certainly colder during glacial periods, but the isotopes are probably not a perfect paleothermometer. Maybe the conditions were different at the moisture sources and maybe preservation of the snow changed. Note that δD and δ18O in ice cores is very strongly correlated with the CO2 concentration in the atmosphere at the time of precip. (this is determined by analyzing tiny gas bubbles in the ice). Clearly, the isotopes are recording climate changes. The timing of these changes (note the rapid deglaciations) can be precisely determined via age dating of the ice using cosmogenic isotope dating (described elsewhere this course). Some ice cores go back very far at low temporal resolution. Other cores, because the snow accumulation rate is greater, have greater temporal resolution but don’t go as far back in time. Ice cores from lower latitudes have been measured as well. However, these ice cores are taken from mountain glaciers- much more difficult! Lonnie Thompson at Ohio State is famous for taking drill rigs via backpack to 6,000 m altitude and above to get drill cores of mountain glaciers. These may tells us more than the high-latitude ice cores because they are more relevant to climate where we live. The Wikipedia entry “Ice Core” is reasonably accurate (not perfect) and very informative.

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Tree Rings: Additional records of climate Researchers have found wood going very far back in time and can date its age with great precision. δ18O of wood varies with δ18O of local precipitation. There are some complicating factors.

Water Isotopes in Hydrology

1) Tracking of storm water in small catchments. Each storm that arrives has the potential to be isotopically different from soil water and groundwater already in the subsurface. If this is the case, we can measure the isotopic composition of stream water over time and see if and when the “new water” from the new storm begins to show up in the stream. The surprising result of studies using this technique is this: Even though stream water begins to rise shortly after a storm begins, most of the water arriving in the stream is “old water” that has been in the catchment since before the new storm.

2) Seasonal Effects- Summer water differs from winter water

Chicago d18O (monthly averages)

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Above is a plot of δ18O in precipitation, averaged monthly, in Chicago for a few years in the 60’s and 70’s. Note the seasonal variation AND the year-to-year variability for any given month. For other global data, go to: http://www-naweb.iaea.org/napc/ih/IHS_resources_gnip.html The data are accessed via the WISER interface: http://www-naweb.iaea.org/napc/ih/IHS_resources_isohis.html#wiser (you may have to register to access the site). The ISOHIS site has animations showing the evolution of isotope values during the annual cycle. Also, John Gibson at Waterloo has a nice web page of GNIP images and animations: http://www.science.uwaterloo.ca/~jjgibson/

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2) Orographic effects- “The altitude effect” and water tracing Mountain ranges usually have isotopically lighter water than surrounding areas.

a) If a storm drops precipitation starting at low elevations, the high elevation precipitation will be isotopically lighter because of the preceding rainout.

b) High elevations are more able to squeeze moisture out of air masses that carry less moisture (because of previous rainout) and are isotopically lighter

c) Lower areas get rain less frequently, and only from air masses that have more moisture and thus have been subjected to less rainout previously.

3) Orographic effects- Paleoaltitude studies e.g., Chamberlain and Poage, 2000, Geology, vol. 28, 115-118. 1) Rain-out effect should increase with time if a mountain range grows in altitude 3) How big is this effect? How much lighter per km elevation?

- Chamberlain and Poage get 2.1‰/km (empirically, I think) - Theoretical predictions are a little difficult; it is easy to predict how much a rising air

mass cools based on adiabatic decompression. If this were the only effect, we would get 7‰ decrease per km- but in reality the latent heat of the water vapor is released to the air as rain/snow forms and the air cools much less than a simple adiabatic model predicts

- Alison Anders uses δ18O of water in mountain belts as an indicator of precip. Dynamics

4) Atmospheric Science - Atmospheric scientists have begin to use water isotopes to better understand how water

vapor is transported and removed from the atmosphere. - Example: If you are interested in how rain is generated by orographic effects, you can

measure the δ18O of the rain and it tells you how much rain-out has occurred prior to the time of sampling and how this changes with time at a given place.

5) Detecting old glacial meltwater Isotopically light water can be found in deep aquifers of upstate New York and the Midwest. Northwestern NY data were collected by Prof. Don Siegel (Syracuse) and his group.

6) Detecting water that is piped in from remote areas

Many cities get their water from far away areas. For example, San Francisco gets its water from the Sierra Nevada, where the water is isotopically lighter than the rain in San Francisco. It is thus possible to detected groundwater leaking from water supply pipes.

7) Paleohydrology Isotope ratios of minerals left behind by ancient waters (typically, calcite, quartz, or clay minerals) reflect the isotopic composition of the parent waters. In many cases, the isotope ratios are easier to interpret than major or trace element chemical data, and may even give information about the temperature of the water.