The O/O Ratio of Atmospheric CO

The oxygen isotopic content of atmospheric CO is mainly determined by interactions between CO and the global reservoirs of liquid water. This follows because direct gas phase interactions of CO with O and HO vapor do not result in O atom exchange [ Francey and Tans, 1987]. When CO dissolves in water, oxygen atoms are exchanged through a mechanism that involves the hydration of dissolved CO to form carbonic acid (HCO). The time scale for dissolved CO to come to isotopic equilibrium with liquid water is the same as the time scale for hydration, i.e., around 30 seconds [ Mills and Urey, 1940]. The O/O ratio of CO in equilibrium with water at 25 C is 1.041 times higher than the O/O ratio of the water. This equilibrium fractionation factor varies slightly with temperature.

Isotopic ratios are generally reported according to

where the value is customarily multiplied by 1000 and expressed in per mille ( ). Bottinga and Craig [1969] suggested using a standard based on CO in equilibrium with the Standard Mean Ocean Water (SMOW) at 25 C [ Craig and Gordon, 1965]. Most recent measurements have been reported relative to the O/O ratio of CO derived from the Pee-Dee Belemnite (PDB) carbonate standard. This standard has an O/O ratio that is 0.22 higher than the Bottinga and Craig standard [ Friedman and O'Neil, 1977].

The O/O ratio of atmospheric CO is primarily determined by exchanges with leaf water, soil water, and surface sea water [ Francey and Tans, 1987; Farquhar et al., 1993]. Oxygen atom exchange with leaf water occurs because a significant fraction of the CO which diffuses into the chloroplasts of leaf cells is not assimilated but diffuses back into the air, and this fraction will have equilibrated isotopically with chloroplast water. Equilibration occurs in spite of the short (< 1 second) residence time of CO in leaves because of the presence of the enzyme carbonic anhydrase, which is concentrated in the chloroplasts of leaf cells and which dramatically speeds up the hydration reaction. Oxygen atom exchange with soil water occurs primarily through CO which is released into the soil by below-ground respiration and which subsequently diffuses into the atmosphere. Oxygen atom exchange with seawater occurs through the exchange of CO molecules across the air--sea interface.

The oxygen isotopic composition of soil water and leaf water vary considerably. Soil water isotopic composition tends to follow the composition of precipitation which is progressively depleted in O relative to seawater towards high latitudes and towards the interior of continents. Chloroplast water, in turn, tends to be enriched in O relative to soil water by evaporation from leaves because HO evaporates preferentially relative to HO. This enrichment of chloroplast water is sensitive to relative humidity and temperature, which can be highly variable [ Dongmann et al., 1974; Förstel, 1978; Zundel et al., 1978].

A global steady-state budget for O of atmospheric CO is shown in Figure 1. This budget uses figures from Farquhar et al. [1993] for fluxes and isotopic exchanges of atmospheric CO with leaf, soil, and sea water. One significant source of uncertainty here is the global average isotopic composition of chloroplast water. Logically, the O of chloroplast water should be intermediate between that of soil water and water at the evaporating surface in the leaves where the maximum isotopic enrichment occurs. A critical question is where does chloroplast water fall in this range. Farquhar et al. [1993] present results based on isotope exchange experiments with several varieties of fruit trees that suggest the isotopic composition of chloroplast water is virtually identical to that of water at the evaporating surfaces in leaves. The budget in Figure 1 is based on this assumption, taking into account the variability of leaf water over the surface of the earth. In contrast, Yakir and coworkers have conducted isotope exchange experiments on sunflowers that indicate that chloroplast water is typically 6 to 10 depleted in O compared to water at the evaporating surface [ Yakir et al., 1993; Yakir et al., 1994]. The difference in O between chloroplasts and evaporation sites probably varies significantly from species to species [ Yakir et al., 1993].

A global model describing oxygen atom exchanges of CO with terrestrial ecosystems has been developed by Farquhar et al. [1993] (see Table 1, Equation G). This model is based on a formulation in which the oxygen--atom exchanges with leaf water are described using an effective fractionation factor (see Table 1) against O on net uptake of CO. The isotopic exchange flux between the atmosphere and leaves is thus obtained by multiplying by net flux of CO into the leaves (basically equal to gross primary production, GPP). The factor is not a true fractionation factor because it depends on the isotopic composition of atmospheric CO. is nevertheless useful because it can be measured in controlled experiments as well as modeled over the surface of the earth [ Farquhar et al., 1993].

The latitudinal distribution of as estimated by Farquhar et al. [1993], is shown in Figure 2. Also shown is the latitudinal variation of CO in equilibrium with surface seawater ( ), the isotopic composition of CO returned to the atmosphere through soils (), the sum , and the annual mean surface values of O of atmospheric CO. Exchanges with leaves and soils tend to drive the local O of atmospheric CO tends towards the sum . This sum tends to decrease towards high latitudes in the northern hemisphere, following the depletion of O/O of precipitation. The latitudinal gradient in can account qualitatively for the latitudinal gradient in O/O of CO that was observed by Francey and Tans [1987], although the actual profile in the air is smoothed by atmospheric mixing.

In addition to exhibiting a gradient with latitude, the O/O ratio of CO is known to undergo a seasonal cycle in the northern hemisphere [ Keeling, 1961; Friedli et al., 1987] with a maximum in early summer and a minimum in early winter. This seasonal cycle probably results mostly from the seasonality of exchanges with terrestrial ecosystems. These exchanges will tend to cause a decrease in O/O of CO during the warmer months when atmospheric CO exchanges most rapidly with leaf and soil reservoirs which are depleted in O/O at middle and high northern latitudes. The ratio will tend to increase during other seasons as a result of transport of higher O/O ratios from more southern latitudes. Other factors, such as seasonal variations in soil water O/O ratios, in leaf water isotopic enrichment, and in the ratio of CO exchange rate with leaves and soils probably also play a role. Modeling this seasonal cycle remains an important area of future work.

What can we learn from measurements of O/O ratios of CO? Farquhar et al. [1993] propose using the measurements to distinguish between CO exchanges with different biomes and between terrestrial ecosystems and the oceans. This application is suggested because of the large variability in the effective discrimination factor between different biomes.

Further insight into what can be learned from measurements of oxygen isotopes of CO can be obtained by the formulation shown in Table 1 (Equation F) which divides the isotopic exchanges with terrestrial ecosystems into three separate terms. The first term is proportional to the gross flux of CO out of leaves which is related to stomatal conductance. The second is proportional to the flux of CO out of soils which, for an ecosystem which is neither gaining or losing carbon, is closely related to ecosystem gross primary production (GPP). The third is proportional to the net flux of CO into the ecosystem, i.e., to net ecosystem production (NEP). The first two terms depend on gross (i.e., two-way) exchanges of CO while the third term depends on the net (one-way) CO exchange. The third term is present because net CO uptake by the ecosystem results in discrimination against O in CO as the CO diffuses into stomata. Under some circumstances, e.g., over a diurnal cycle, the instantaneous value of this third term may be comparable in magnitude to the first two terms. On a time averaged basis, however, this term will be relatively unimportant because NEP is generally at least an order of magnitude smaller than the flux of CO out of soils or leaves. (Figure 1 was drawn assuming this term is zero). What this means is that O of CO is mainly sensitive to gross rather than net exchanges [see also Yakir et al., 1993]. The unique value of the O measurements therefore lies in providing new information on rates of gross primary production and stomatal conductance, as these exchanges can produce large variations in O without producing variations atmospheric CO concentration or carbon isotopes of CO.

At present, several research programs are engaged in measuring the O/O ratio of atmospheric CO. These measurements will be useful for validating global-scale numerical models including physiologically based exchanges of HO and CO with leaves and soils. The O/O measurements can be expected to provide information on rates of gross primary production and stomatal conductance integrated over large spatial scales and in variations in these quantities in response to climate change, increasing atmospheric CO, or other global variables.