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Simulation of Oxygen Isotopes in a Global Ocean Model A. Paul*, S. Mulitza, J. Pätzold and T. Wolff

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Simulation of Oxygen Isotopes in a Global Ocean Model A. Paul*, S. Mulitza, J. Pätzold and T. Wolff Simulation of Oxygen Isotopes in a Global Ocean Model A. Paul*, S. Mulitza, J. Pätzold and T. Wolff Universität Bremen, Fachbereich Geowissenschaften, Postfach 33 04 40, D-28334 Bremen, Germany *corresponding author (e-mail): [email protected] δ18 Abstract: We incorporate the oxygen isotope composition of seawater Ow into a global ocean model that is based on the Modular Ocean Model (MOM, version 2) of the Geophysical Fluid Dynamics Laboratory (GFDL). In a first experiment, this model is run to equilibrium to simulate the present-day ocean; in a second experiment, the oxygen isotope composition of Antarctic Surface Water (AASW) is set to a constant value to indirectly account for the effect of sea-ice. We check the δ18 depth distribution of Ow against observations. Furthermore, we computed the equilibrium δ18 fractionation of the oxygen isotope composition of calcite Oc from a paleotemperature equation δ18 δ18 and compared it with benthic foraminiferal O. The simulated Ow distribution compares fairly δ18 well with the GEOSECS data. We show that the Ow values can be used to characterize different δ18 water masses. However, a warm bias of the global ocean model yields Oc values that are too light by about 0.5 ‰ above 2 km depth and exhibit a false vertical gradient below 2 km depth. Our ultimate goal is to interpret the wealth of foraminiferal δ18O data in terms of water mass changes in the paleocean, e.g. at the Last Glacial Maximum (LGM). This requires the warm bias of the global ocean model to be corrected. Furthermore the model must probably be coupled to simple atmosphere and δ18 sea-ice models such that neither sea-surface salinity (SSS) nor surface Ow need to be prescribed δ18 and the use of present-day Ow-salinity relationships can be avoided. Introduction natural assembly of all isotopic species. Therefore, Isotopes in the Hydrological Cycle water vapor evaporating from the sea surface is Measurements of the isotopic composition of wa- depleted of heavy isotopes relative to ocean water, ter at the various stages of the hydrological cycle while rain precipitating from a cloud is enriched has enabled the identification of different water relative to the cloud moisture. Hence whenever a masses and the investigation of their interrelation- water sample undergoes a phase transition (e.g. ships; measurements of the variations of the iso- evaporation or condensation), temperature-depend- topic composition of proxy materials in climate ar- ent isotope fractionation occurs (Fig. 1). In the case chives such as ice and sediment cores has become of tritium, this effect is generally masked by mixing a most useful tool in paleoclimate reconstructions water from different sources (Jouzel 1986). But in (Gat 1996; Schotterer et al. 1996). the case of deuterium and oxygen-18, the isotopic The isotopic species of the water molecule that composition of precipitation shows a large variation are of most interest to the geosciences are with latitude, height and continentality. The deple- 1 2 16 1 18 1 3 16 H H O, H2 O and H H O. A water molecule tion of precipitation in deuterium and oxygen-18 is containing the radioactive isotope 3H (tritium) or most pronounced in the polar regions. The variation either of the stable isotopes 2H (deuterium) and 18O in the isotopic composition of seawater is compara- (oxygen-18) is heavier than the water molecule tively small and mainly determined by freshwater 1H2H16O that is by far the most abundant in the input and mixing between water masses. From FISCHER G, WEFER G (eds), 1999, Use of Proxies in Paleoceanography: Examples from the South Atlantic. Springer-Verlag Berlin Heidelberg, pp 655-686 656 Paul et al. The stable isotopes deuterium and oxygen-18 The δ value that is commonly used to express have been incorporated into a number of atmos- the isotopic composition of a water sample is de- pheric models, e.g. the LMD model (Joussaume et fined by al. 1984; Joussaume and Jouzel 1993), the NASA/ GISS model (Jouzel et al. 1987), the ECHAM δ = (R / R −1)⋅1000‰, (1) model (Hoffmann and Heimann 1993; Hoffmann sample standard 1995) and the GENESIS model (Mathieu 1996). As these four models account for the various where Rsample and Rstandard refer to the isotope 2 1 18 16 fractionation processes that occur at the sea sur- ratio H/ H or O/ O in the water sample and in δ face and all succeeding stages of the hydrological the standard. Thus positive values indicate an cycle, they provide a tool for calculating the isotopic enrichment of the heavy isotopic species relative δ content of precipitation δP as well as of evapora- to the standard, and negative values indicate their tion δE (Juillet-Leclerc et al. 1997). Prior to their depletion. The generally used standard is the Stand- development there was no method to estimate δE ard Mean Ocean Water (SMOW, Craig 1961) or, independently (Craig and Gordon 1965). more recently, the Vienna Standard Mean Ocean Water (V-SMOW, Gonfiantini 1978). Fig. 1. Sketch of the global δ18O cycle. In our simulations, only the oceanic part has been considered so far. Simulation of Oxygen Isotopes in a Global Ocean Model 657 The focus of this paper is on oxygen isotopes, of interest (Weiss et al. 1979; Jacobs et al. 1985; because oxygen-18 variations are directly related Bauch 1995; Toggweiler and Samuels 1995). to paleotemperature studies. It is in fact for this To give an explicit example, we reproduced the reason that most of the oceanographic work to date detailed freshwater budget for the Ross Sea conti- has concentrated on oxygen-18 and only few deu- nental shelf constructed by Jacobs et al. (1985). terium measurements have been performed (Craig From the difference between the δ18O content and Gordon 1965). of shelf water and its source they estimated that 36 cm a-1 of glacial meltwater is introduced to the shelf. They found that surface and deep waters Applications in Present-Day Oceanography move onto the shelf with an average δ18O content The distributions of both salinity S and oxygen iso- of -0.22 ‰. On the shelf, its δ18O content will be δ18 δ tope composition of seawater Ow or simply w lowered by marine precipitation, glacial meltwater are mainly controlled by the same two processes and the net freezing of sea ice, such that the aver- δ δ18 precipitation P and evaporation E. Locally S and w age O content of shelf water becomes -0.42 ‰: of surface waters are linearly related. However, the (1) Marine precipitation at a rate of 15 cm a-1 with δ δ18 δ18 slope of the w-S relationship varies between 0.1 an O content of -16.5 ‰ will lower the O for tropical surface waters and 1 for high-latitude content of shelf water by 0.005 ‰ a-1. (2) Further surface waters. This reflects the so-called tempera- depletion of 0.025 ‰ a-1 will result from the addi- ture effect: in high latitudes, precipitation occurs at tion of glacial meltwater at a rate of 36 cm a-1 with lower temperatures and is more enriched in oxy- an δ18O content of -36 ‰. (3) The decrease in δ gen-18 than in low latitudes, which leads to a strong shelf water w from sea ice freezing will be only depletion of the remaining cloud moisture. A slope 0.002 ‰ a-1, due to the near-zero slope of the freez- of 1 is observed for East Greenland surface wa- ing line. The net change in δ18O by -0.20 ‰ a-1 and ters, which are diluted by adding meltwater from the glacial addition of meltwater at a rate of 36 cm the permanent ice cap. Freshwater input from riv- a-1 require a shelf water residence time of 6.2 a. δ ers can also modify the w-S relationship. Hence S An important application of a detailed freshwa- δ and w label the surface waters in different ways. ter budget is the study of the role of sea ice in the Mixing of water masses that originate from the sea formation of Antarctic Bottom Water (AABW) surface creates the characteristic vertical sections (Toggweiler and Samuels 1995; Stössel et al. 1996). known from GEOSECS (Birchfield 1987; Östlund In the example of Jacobs et al. (1985), the onshelf et al. 1987). flow has an average salinity of 34.455 and the The isotopic composition of sea water is also offshelf flow has an average salinity of 34.589, affected by the formation of sea ice, but the which leaves a salinity enrichment of 0.134 units. fractionation effect is so small that the freezing This salinity enrichment corresponds to the brine process leads to an increase in salinity, with essen- rejected during the formation of sea ice at a rate of δ -1 tially no observable influence on w. This gives a 37 cm a , with a residual salinity of 4.1, from Ant- δ near-zero slope in a w-S diagram (Craig and arctic Surface Water (AASW) with a salinity Gordon 1965). As a result, the surface waters of of 34.1. Furthermore salt must be drained from the Southern Ocean show a wide range of salinities, 58 cm a-1 of sea ice to balance 15 cm a-1 of marine δ -1 although w is nearly constant. For Antarctic Sur- precipitation and 36 cm a of glacial meltwater. face Water (AASW), Jacobs et al. (1985) report a Toggweiler and Samuels (1995) pointed out that on δ w value of -0.31 ‰ with a standard deviation of the one hand brine rejection causes a net salinity 0.08 ‰.
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