Oceanography 540--Marine Geological Processes--Autumn Quarter 2002

Oxygen and Carbon Isotope as Paleotracers

Oxygen Isotopes

Oxygen has three different stable isotopes: Molecules containing atoms of differing mass behave differently from one another:
  1. in chemical reactions: the molecules have different chemical potentials and so have different equilibrium constants
  2. during mass transport and in rate-limited chemical reactions: the molecules have different masses--the light molecule is transported or reacts more quickly
For example, at equilibrium:

Eq 1: eq 21-1


Eq 2: eq 21-2

If a fractionation factor, alpha, is defined as the ratio of the isotopic ratios in two phases, it can be shown that this fractionation factor is equal to K:

Eq 3: eq 21-3

This equilibrium fractionation is usually expressed in a different way, as the "del value" for the fractionation (upper case Delta):

Eq 4: eq 21-4

which is usually expressed per mil, ‰

Similarly measurements of the isotopic ratio are generally expressed in a parallel del notation (lower case delta):

Eq 5: eq 21-5

also expressed per mil, ‰.

For oxygen isotopic measurements of water the standard is standard mean ocean water, SMOW, while for carbonates either SMOW or fossil belemnite from the Pee Dee formation, PDB, is used as the standard.

A consequence of the equilibrium relationships is that if two materials equilibrate then

Eq 6: eq 21-6

To apply this relationship in interpreting geological record, we are concerned with four possible effects on the temporal variability:

  1. the isotopic composition of seawater
  2. the temperature of the water in which the calcium carbonate was deposited
  3. whether equilibrium is established
  4. whether calcium carbonate acts as a closed system once deposited
The isotopic composition of liquid water responds to evaporation and precipitation. There are combined effects of differences in vapor pressure of the light and heavy molecules, i.e., in the equilibrium between vapor and liquid, and in the mass-dependent transport across the air-sea interface. The vapor pressure effect is about -8.5 to -10 ‰, while the mass transport effect is an additional -2 to -4 ‰, i.e., the overall shifts are between -10.5 and -14 ‰, the vapor being isotopically lighter than the water. In approximate terms:

deltasub v-deltasub l=-12 ‰=Deltasub vsub l
Deltasub vsub l=alphasub vsub l-1
-12 ‰=-.012=alphasub vsub l-1
alphasub vsub l=1-.012=.988

As water vapor condenses to form clouds there is an opposite effect, the first precipitation is isotopically heavier, leaving lighter water vapor behind and alphasub lsub v~1.012.

The marine atmosphere is fairly well buffered by the isotopic composition of the surface ocean (see below for the range of variation in the surface ocean). Once vapor is transported over the continents, a Rayleigh condensation begins. This condensation generally occurs at lower temperature than did the evaporation which provided mositure to the cloud and at lower temperature the fractionation between the two isotopes is greater. In terms of the fraction of residual vapor, f:

Eq 7: eq 21-7

Rsub v is the 18/16 ratio in the vapor and the superscript * denotes the initial vapor. Consider a cloud containing vapor with delta^1^8O=-10 losing moisture as precipitation according to equation 20-7 (As the cloud cools, the value of alpha will progressively decrease; this is not accounted for here). The progression and the liquid and vapor composition are shown here:


Rayleigh condensation

Figure 1


During glacial times storage of the isotopically light water produced at moderate (f<0.01) on the continents leaves behind an isotopically heavier ocean.

By examining the isotopic composition of benthic forams, the effects of changing temperature and the change in the isotopic composition of seawater can be separated. We look at benthic forams because in the deep ocean the temperature decrease during glacial times is constrained by the freeezing point of seawater. There is a 1.9 ‰ difference in oxygen isotopic composition between glacial and interglacial times with the glacial value being higher. The maximum temperature shift in the deep ocean is of order of the present day 1°C to the freezing point of about -1.8°C. This 2.8°C temperature shift can lead to an isotopic shift of no more than 2.8 x 0.22 ‰/°C or 0.6 ‰. A shift of at least 1.3 ‰ remains to be explained by the change in the isotopic composition of the water, and a still greater shift if the temperature difference in deep water is less than 2.8°C.

Forming a mass balance in terms of the volume of the present day ocean, Vsub osub c, and the glacial storage of ice, Vsub isub csub e, and assuming that the mean isotopic composition of glacial ice is -35 ‰ (consistent with ice core measurements):

Eq 8: eq 21-8

Dividing through by Vsub osub c and letting x=Vsub isub csub e/Vsub osub c:

Eq 9: eq 21-9

This equation can be solved to find that x~.035, corresponding to .035 x 4000 m ~ 140 m of sea level change. This decrease in volume by evaporation leads to an average increase in ocean salinity to 1.035 x 35 ‰=36.2 ‰.

In summary: the oxygen isotopic signal in benthic forams is primarily an ice volume signal (and secondarily a temperature signal).

Another contributing factor to variations in seawater oxygen isotopic composition is the internal distribution of oxygen isotopes in the ocean. In the present day ocean:



Figure 2


NADW maintains the signature of surface waters because of its mode of formation, by cooling-induced convection, whereas AABW does not because of the salinity increase brought about by salt expelled during formation of sea ice. The isotopic shift experienced by benthic forams can thus also be influenced by the history of formation of deep ocean waters.

Do planktonic forams have a temperature signal in addition to the ice volume signal? The ice volume effect dominates the overall signal, but at particular locations an additional component can be recognized amounting to about 2°C cooler temperatures in surface waters during glacial times. However analysis of this signal depends on correctly estimating the surface water oxygen isotope distribution which adds considerable error.

Carbon Isotopes

The important stable isotopes of carbon are: Carbon isotopes are expressed relative to the PDB standard, belemnite carbonate from the Peedee formation, South Carolina. (Oxygen isotopic composition of oxygen in carbonate may also be expressed on this scale; it is equivalent to +30.6‰ on the SMOW scale).

delta^1^3C provides a measure of oxidation of organic carbon. There is little fractionation of carbon isotopes between COsub 2 dissolved in the water and solid CaCOsub 3:

Eq 10: eq 21-10

There is a much larger fractionation when organic carbon is produced:

Eq 11: eq 21-11

As deep ocean waters age, the evolve from being nutrient poor, oxygen saturated, CO2 depleted as oxidation of organic carbon and dissolution of calcium carbonate occur. For all practical purposes, only oxidation of organic carbon is important in the isotopic balance because of the very small fractionation between dissolved carbon dioxide and calcium carbonate. The concentration of some nutrient, say phosphate, can be used as a measure of the extent of oxidation of organic carbon. Thus in the present day ocean, the isotopic composition of COsub 2 can be related to the phosphate concentration through the mass balance:

Eq 12: eq 21-12

Eq 13: eq 21-13

Eq 14: eq 21-14

In the present day ocean this linear correspondence of isotopic composition of dissolved inorganic carbon and the phosphate concentration is observed with the range being about 2.6‰ (-0.85 x 3 µM phosphate).

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