Oxygen isotopes
H2O is composed of both H216O and H218O
⎛ H 2 18O ⎞
⎛ H 2 18O ⎞
⎜ 16 ⎟ − ⎜ 16 ⎟
⎝ H 2 O ⎠ ice ⎝ H 2 O ⎠ SMOW
18
δ Oice =
*1000
‰
⎛ H 2 18O ⎞
⎜ 16 ⎟
⎝ H 2 O ⎠ SMOW
Hydrogen isotopes
H2O also contains heavy hydrogen (2H), also known as
deuterium (D).
⎛ D ⎞
⎛ D ⎞
⎜ ⎟ − ⎜ ⎟
⎝ H ⎠ ice ⎝ H ⎠ SMOW
δDice =
*1000 ‰
⎛ D ⎞
⎜ ⎟
⎝ H ⎠ SMOW
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Relationship between K and alpha for water
evaporation
H 2O(l) + H 218O(g) ⇔ H 218O(l) + H 2O(g)
ΔG = −RT ln(K)
K≡
[H 218O(l)][H 2O(g)]
K =α =
(
([H 218O(g)][H 2O(l)]
)liquid
H 218O /H 2O
(
)gas
H 218O /H 2O
Take note that aeq will not in general be equal to KEQ because the ratio of isotope abundance will not be the
same As the ratio of molecular abundances of the isotope-containing molecules. For 18-O/16-O in
evaporating water, it happens to be simple!, but this is not usually the case.
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ΔGrxn = −RT ln(K)
= 22 J at 25 C
K = 0.991
α EQ = K EQ = 1.009
Take note that aeq will not in general be equal to KEQ
because the ratio of isotope abundance will not be the same
As the ratio of molecular abundances of the isotope-containing
Molecules. For 18-O in water, it happens to be simple!
H
(
α=
18
2
gas
= p16 / p18
€
gas
(H
HO
(
=
2
)
O /H 2O
/H 2Oliquid )
(
18
2
)
O /H 2O
18
liquid
18
H 2 Ogas /H 2 Oliquid
)
Cartoon of water transport and isotope fractionation from
ocean to ice sheet. Note that in pure Rayleigh
fractionation, there is no re-evaporation from the ocean surface
(shown the the up/down curved arrows)
Raleigh distillation model
We can track the progression
of the vapor-rainfall if we know:
1.  the initial isotopic ratio of the vapor
2.  the fraction of vapor remaining
RV
= f α −1
RV0
δ18O (‰
liquid
where RV is the isotopic ratio of the vapor
RV0 is the initial isotopic ratio of the vapor
ƒ is the fraction of vapor remaining
α  is the fractionation factor
vapor
After Dansgaard, 1964
Why is the fractionation increasing
(9, 10, 11) in this figure?
Temperature vs. saturation vapor pressure of water
Fractionation of vapor remaining as a function
on temperature in a cooling, saturated air mass, initially at 25°C
‘fraction of original water vapor remaining’ = saturation vapor pressure (at temperature T)/
saturation vapor pressure (at temperature Tinitial)
Water isotope ratio (δ18O) in the vapor phase, in a saturated
Cooling air mass, assuming that the initial value is -9‰,
and that alphaequilibrium is constant at 1.009
Relationship between T and δ18O, in a simple (air saturated,
alpha = constant) Rayleigh fractionation model
0
18
δ O (‰)
-5
-10
δ18O = 0.58T - 23
r2 = 0.9995
-15
-20
-25
0
10
20
temperature (°C)
30
Relationship between snow accumulation and temperature (estimated from d18O) in Greenland ice cores. Using ClausiusClapeyron equation to estimate temperature change from the observed accumulation change reproduces the isotope-inferred
temperature change very nicely. This appears to validate the simple Rayleigh fractionation model for Greenland; note
however (see caption) that the temperature vs. d18O slope is much lower than expected from that simple model. This is
though to be due to precipitation seasonality bias (during cold climate periods, the loss of winter snowfall is greater than the
loss of summer snowfall, so the mean annual isotope ratios tend to look ‘too warm’ (they are biased towards summer).
Temperature effect on the δ18O of precipitation
holds for both spatial T variability
and temporal variability
The traditional assumption is that the spatial slope between
T and d18O is a good estimate of the temporal slope.
This is an assumption that appears to hold on interannual
timescales, but generally not on longer timescales, due to
changes in seasonality of precipitation, changes in moisture
source trajectories, etc. Refer to Jouzel et al., 1997 for
details.
Rozanski, 1993
Law Dome (Antarctica) Stable Isotope Record
Corresponds to about 15 K Law Dome (Antarc8ca) Stable Isotope Record Law Dome is a very high accumula8on site (1 m /year of ice) and therefore preserves water isotope variability beau8fully Seasonal variations in d18O are not necessarily a temperature signal!
The temperature/d18O relationship holds only at mid-high latitudes.
What is happening at
higher temperatures?
In the tropics, the
‘amount effect’
dominates
Rozanski, 1993
The so-called “amount” effect: more rain, lighter δ18O
NOTE: mostly in tropics (<30° N and S), where “deep convection” takes place
Empirical relationship
It would be difficult to explain
a vapor source at +1‰, when the tropical oceans are ~0‰.
Dansgaard, 1964
Rozanski, 1993
Thought to be linked to increased evaporation of raindrop in dry,
under-saturated environment…
(raindrop is enriched as it falls
from the sky)
Mechanism still unknown – needs
atmospheric modeling work
The “Global Meteoric Water Line” – what happens to δ18O happens to δD, but with a different α
annual mean δD vs. δ18O of precipitation
But month-to-month variations
at a given site fall off this line – “deuterium excess”
d = δ D − 8* δ 18O
δ D = 8* δ 18O + 10
Craig, 1961
Rozanski, 1993
Why don’t all waters fall on the GMWL?
Or…. why do different “source” waters have different ‘deuterium excess’ values?
- evaporation not purely equilibrium process
- what other type of fractionation
is involved?
1-3km
Fact: water vapor above the ocean is -13‰ in δ18O, not the -9.2‰ expected from equilibrium fractionation. Why? Water
Given the potential for complicated boundary layer physics, it’s a wonder that the
GMWL exists at all!
Surface Water Salinity-δ18O relationship - general
δ 18O = 0.45* S − 15.5
Global precipitation
So δ18O of surface waters, like salinity,
is also correlated to evaporation – precipitation.
Intermediate complexity models:
-50
Vostok
δ18O (‰)
-52
-54
-56
-58
-60
-62
0
50
100
150
Age (thousands of years before present)
200