“Water Isotopes:” δ18O and δD in “Water Isotopes:” δ18O and δD in Hydrology Reading:...

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Transcript of “Water Isotopes:” δ18O and δD in “Water Isotopes:” δ18O and δD in Hydrology Reading:...

  • 5/10/09

    Water Isotopes: 18O and D in Hydrology Reading:

    Clark and Fritz, Chs. 2 - 4 Additional resources: Isotope Tracers in Catchment Hydrology, Kendall and McDonnell, eds. Elsevier, 1998.

    A very excellent resource. Bill Whites lectures #26 and 27 from Cornells Isotope geochem course Motivation: - 18O and D vary in nature- heavy vs. light isotope fractionation in several common processes - These variations provide opportunities to trace water; its nice that the actual constituents of the

    water are what we measure so we are tracing water itself and not some dissolved constituent that may or may not follow the water well.

    A note on the standards used to define 18O and D Ocean water is nominally the zero point for 18O and D scales The names used for the international standards are SMOW or VSMOW

    SMOW (standard mean ocean water) was real ocean water; supply is all used up VSMOW is a replacement used now

    Thus, a measurement result of 18OVSMOW = +1.0 means the samples 18O/16O ratio is one part per thousand greater than that of VSMOW, or ocean water. Actually, ocean water varies a bit with location and time, so SMOW was a bit arbitratry 18O measurements of carbonate materials are compared to the PDB (now VPDB) standard, which has a greater 18O/16O ratio by about 31 .

    Fractionation of O and H: equilibrium vs. 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 fractionations-

    At 0C: 1000ln ( liquid - vapor) = 11 for oxygen isotopes = 99 for hydrogen isotopes At 100C: = 3.2 for oxygen isotopes = 29 for hydrogen isotopes (From Faure, 1986)

  • 5/10/09

    The Hydrologic Cycle: Systematics of 18O and D in water vapor in Earths Atmosphere Rayleigh model gives a rough approximation of reality on earth: Dansgaard (1964) published a paper with:

    Measurements of 18O plotted versus temperature. Rayleigh models of 18O versus temperature

    The classic Rayleigh model: 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 you cool a packet of air, it drops moisture as rain or snow 5) The transport of moisture poleward thus involves a distillation process 6) The fraction of the original moisture remaining should be a function of temperature 7) Using a Rayleigh fractionation model, we can then link temperature to 18O It is interesting that the 18O is a nearly linear function of T. How can this be, when the Rayleigh relationship is such as non-linear function of amount remaining?

    By chance (or maybe theres some hidden connection in the physics), the fraction of water vapor remaining is close to: f AeBT

    Whereas the Rayleigh model is close to 18O - 18O(initial) = ln(f) So the ln(x) and ex functions sort of cancel out to give a nearly linear result

    The original figure from Dansgaards 1964 paper. Even though the correlation between T and d18O looks very strong, notes two things: 1) Most of the line is defined by data from very cold places. Plots of data from places where people actually live show strong scatter. 2) Some points (especially islands) fall pretty far from the line.

  • 5/10/09 Reality is more complicated 1) Poleward vapor transport is more like diffusion than advection. Rayleigh models assume closed systems (e.g., a packet of air moving from high T to low T; no mixing with other air masses; rain/snow allowed to leave the air, no other gain/losses)

    This assume water moves poleward mostly via advection But in reality, theres much mixing between air masses

    An alternative model: Transport is mostly by random mixing of air masses, which is NOT advection. It is much closer to diffusion Eddy diffusion is the term used to describe this process Eddy diffusion model gives differen slope to the isotope/temperature plot. Hendricks et al. (2000) [Global Biogeochem. Cycles, Vol. 14, pp. 851-861]

    2) Air masses can gain H2O from local sources. This will tend to alter the 18O of the air mass, unless by chance the new vapor matches the exisiting vapor.

    - oceans (after the air leaves the tropics) Local ocean water is about 0; vapor derived only from local ocean water

    would give rain close to 0 even at low T (in reality the vapor is not all local) so coastal and island areas may have 18O vs. T relationships that differ from

    the normal one (18O shifted upward; see island on plot above) - lakes, rivers, soils, too

    3) Orographic effects: 1. Mountain ranges have strong rainout 2. As a results 18O is more negative on the downwind side 3. Also, the air is warm because of the latent heat it gained as rainout occurs over the

    mountains (e.g., Nevada is hot in part because of strong rainout over the Sierra Nevada)

    4. This alters the classic delta vs. T relationship 5. How big is this effect? 1.0 to 3.1 per km of elevation increase

    This has been used to find elevation of mountain ranges in the past clay minerals in the downwind are record rainout changes as the

    mountains grow See: Chamberlain and Poage, 2000, Geology 28, 115-118

    4) The amount effect: altitude gradients in 18O, and re-equilibration or evaporation of rain as it falls It has been observed that, all other things being equal, more intense rain means lighter isotopic composition. This is attributed to a few different effects. Fortunately, they all tend to make heavier rains lighter: - Light rains can re-evaporate in the lower atmosphere, making them heavier - Water vapor is lighter with increasing altitude; rain starting high up is thus relatively light. But if

    it falls slowly (smaller drops; lighter rain) it requilibrates with vapor at lower altitude and becomes heavier. Intense rain means large drops that do not re-equilibrate and thus retain lighter isotopic composition from high altitude.

    - Rain moving from high levels to lower makes the vapor at lower altitude lighter, which in turn allows rain to stay lighter as it partially re-equilibrates with lower air. This effect is stronger for heavier rains

    - Heavy rains are more likely to wring large amounts of moisture from the air, to the point where the vapor remaining toward the end is shifted to more negative values.

    - Reference: Amount effect of water isotopes and quantitative analysis of post-condensation processes. J.-E. Lee and I. Fung, Hydrol. Process. (2007)

    - The amount effect is not seen with snow; re-requilibration and sublimation are minor

  • 5/10/09 End result of all these complications: - DO USE your general knowledge of how temperature and other factors affect the isotopic

    composition of precipitation to predict general patterns you might use to your advantage. - But dont use the global relationship between T and 18O blindly; there may be deviations for

    various places, times, and precipitation events. There are often differences with time in a single storm.

    Relationship between D and 18O on earth: Meteoric water lines and how to detect evaporation or water-rock interaction.












    -16 -12 -8 -4 0 4

    delta 18O


    a D

    1) Evaporation and precipitation effects combine to give us the observed array of precipitation

    on a plot of D vs. 18O a) seawater plots at the origin b) Evaporation of that water produces vapor plotting at roughly (-11, -90) c) The first precipitation from this vapor would be fractionated back toward lower

    values d) However, precipitation is not an exact reversal of evaporation, and the first rain

    formed from this vapor plots above the seawater point. e) Subsequent precipitation causes both O and H to become isotopically heavier, with a

    slope of about 8 (varies- see below) i. Example: When 86% of the vapor has rained out, 18O of the next rain will be

    about -20.

    2) The solid line in plot above gives the approximate compositions of rain/snow on earth. This is known as the Global Meteoric Water Line (GMWL). D =(7.96) 18O + (8.86) (Ronzanski et al., 1993)

    3) Evaporation from lake/river/puddle/ocean surfaces. 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.

    a) Why is the slope different? Kinetic fractionation during evaporation is a very different process from equilibrium fractionation during precipitation. Recall how kinetic and equilibrium effects are controlled by different things. Because of this difference, the slope (size of hydrogen isotopic fractionation relative to that of oxygen) is different.

    Plot of D vs. 18O: Close linear relationship between 18O and D, 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. Vapor derived from seawater plots near (-12, -80); exact values depends on close to Precipitation samples from cooler places plot along the GMWL; colder places are farther to the lower left. The horizontal arrow