Encart III.1. Le mod•le de surface ˆ une boite
III.3. Les rŽsultats de notre simulation globale
For example (Broecker, 1986), the difference of salinity (and of isotope) between surface and deep waters has, for a given ocean, a different value for Present-day and Last Glacial Maximum (21 000 years Before Present) climates. Recently, Rohling and Bigg (1998) critically assessed the reconstruction of paleosalinity based on the assumption of a constant and linear oxygen 18- salinity relationship. Instead, they concluded that both the spatial and temporal variabilities of this relationship need to be constrained.
One way to approach these problems is to model the oxygen 18 distribution in the ocean and its relationship with salinity for different climates. This strategy, parallel to that followed to interpret polar and continental paleoprecipitation (Jouzel et al., 1999), has been suggested by various authors (Juillet-Leclerc et al., 1997, Rohling and Bigg, 1998). The most promising results have been obtained by Schmidt (1999b), thereafter SC99 using the GISS (Goddard Institute for Space Studies) oceanic general circulation model (GCM) forced by isotopic fluxes derived from the GISS atmospheric GCM (Jouzel et al., 1987). An alternate approach, by assimilating observed surface isotopic content, has also been successfully applied (Wadley and Bigg, 2000, Paul et al., 1999). But this method remains sensitive to the number and distribution of observations, and precludes any paleoclimatic developpment.
Hereafter, we follow a similar modelling approach to SC99, with a different oceanic GCM. Our work is specific in that 1) we use consistent hydrologic fluxes predicted by an atmospheric GCM implemented with water isotopes (whereas SC99 uses water fluxes -salinity- calculated by the GISS coupled model), 2) we perform long experiments (> 2000 years) which allows us to investigate oxygen 18 distributions in surface to deep waters (instead runs presented in SC99 cover less than 100 years which limits interpretation to surface characteristics), and 3) the ocean starts with an homogeneous isotopic composition (the oceanic composition in SC99 is instead initialized to observed values), which is somewhat an insurance that realistic mechanisms are simulated by the model. Based on a present-day simulation, we mainly discuss the consequence of the ocean dynamics on the isotope distribution, compared to the 'static' box model of Craig and Gordon (1965), to extend the work of Juillet-Leclerc et al.
(1997). This simulation leads to substantial improvement compared to this box model, and reasonably captures the main observed characteristics of the oxygen 18 distribution and of the oxygen 18 - salinity relationship.
III.3.2. The atmospheric fluxes and the oceanic tracers.
Atmospheric hydrologic fluxes (evaporation and precipitation, and their isotopic content) are taken from the isotopic version of the NASA/GISS atmospheric GCM (8¡x10¡) developped by Jouzel et al. (1987). This ensures the consistency between water (evaporation and precipitation) and isotopic fluxes, of primary importance here since their balance generate our tracers, as discussed in Juillet-Leclerc et al. (1997). Continental runoff (including iceberg
discharge) is taken from a work of Russell and Miller (1990) who modelled the continental drainage using the GISS GCM. The annual isotopic content of the river discharges is calculated by combining the GISS isotopic outputs with the drainage map of Russell and Miller.
Hydrological fluxes due to sea ice formation and melting are not taken into account because of their complex seasonal variations. Although they play an important role in forming bottom waters, they probably have a limited effect on both salinity and oxygen 18 characteristics of these waters in the real world (Toggweiler and Samuels, 1995).
The oxygen 18 flux to the oceanic surface is expressed as :
F
18O= E
.(R
S- R
E) - P . (R
S- R
P) - R . (R
S- R
R)
where
R
S is the isotopic ratio of the oceanic Surface,R
E ,R
P andR
R are the isotopic ratios of Evaporation(E)
, Precipitation(P)
and Runoff(R)
.Since an important issue of this work is the oxygen 18 - salinity relationship, it is necessary to simulate a salinity tracer consistent with the oxygen 18 one, in that both originate from the same atmospheric fluxes and undergoe the same oceanic mixing. Therefore, we define another tracer by its atmospheric flux
F
S, similar toF
18O above but with Rs being the surface salinity and others R equal to zero. This passive tracer, called ÔsalinityÕ in the following, has no dynamical effect since we calculate it offline, but it is transported by the model in the same way than the dynamical salinity is. The basic understanding of these atmospheric fluxes,F
18O andF
S , is that evaporation tends to enrich the surface in salt as well as in oxygen 18 with respect to oxygen 16, contrary to precipitation and runoff.Both fluxes,
F
18O andF
S , are prescribed monthly to a tracer version of the OPA oceanic GCM of the Laboratoire dÕOcŽanographie Dynamique et de Climatologie (LODyC, Paris-France), a primitive equation model described by Madec et al. (1998), run in 'robust- diagnostic' mode. Atmosphere-ocean interaction is described in detail by a fine subsurface resolution (10 levels in the first 100m) and a turbulent kinetic energy parameterization of the vertical diffusion. Although the degraded version used here has a medium horizontal 92x77 resolution, it benefits from its parent (180x150) dynamics averaged on its own resolution, and reasonably simulates distribution of geochemical tracers like 14C (Aumont et al., 1998). This seasonal ocean dynamics serves here to mix 'offline' the tracer variations generated by the atmospheric fluxes. Note that this dynamics is not consistent with our water atmospheric flux (F
S) since it is calculated by restoring to observed temperature and salinity, and not by flux forcing.The simulation begins with an homogeneous oceanic composition of 34.6 permil for the salinity and V-SMOW (Vienna-Standard Mean Ocean Water with an H218O/H216O ratio, RSMOW, of 2005.2*10-6) for the oxygen isotopic ratio. The tracer model is first run, for a spin- up period of 2000 years, with a lower resolution version (47x39) but with the same dynamics degraded to this resolution. Outputs from this simulation serve as initial conditions for the 92x77 version run, which lasts 200 years (this acceleration procedure is described in Aumont et
al., 1998). At this time, because the atmospheric fluxes are not corrected and thus not exactly balanced, a slight drift of the salt and isotope masses remains, but it is considered as negligeable for our first order study : over the last 100 years of simulation, both salinity and oxygen isotopic ratio increase by only 0.004 permil in the first layer. Averaged over the whole oceanic surface, the precipitation flux is about 1.22 m/yr, evaporation 1.28 m/yr and runoff 0.09 m/yr, so that the global water flux imbalance is +.03 m/yr. The time step for the tracer transport is one day.
To illustrate how the atmospheric fluxes,
F
18O andF
S , force the oceanic surface composition, isotopic content and salinity are first estimated without oceanic horizontal advection. This estimation is established from the conceptual 2-box model described by Craig and Gordon (1965) where the surface box composition is forced by the atmospheric fluxes and buffered by exchange with an underneath box reservoir. The model parameters are derived from an empirical correlation between the GISS AGCM isotopic flux and GEOSECS observed surface d18O, computed by Juillet-Leclerc et al. (1997). They infer a vertical mixing flux of 4.8 10-4 mm/s, and a deep water d18O of 0.22 permil. To compute the salinity, we use a deep value of 34.6 permil. This box model is expected to represent an ideal ocean where the prescribed atmospheric fluxes are balanced only by vertical mixing, so that the simulated tracer fields (Figures III.10a and b with the isotopic content expressed as the deviation d with respect to the V-SMOW, in permil : d = (R/RSMOW - 1)*1000) are representative of the atmospheric forcing.The main enrichment zones are centered above subtropics, along both basin sides in the Pacific, along the south-east side and north-west side (Gulf of Mexico) in the Atlantic. The Indian and western Pacific equatorial zones show up a strong depletion due to precipitation, not seen in the Atlantic. In subglacial latitudes, the limit between enrichment and depletion lies around 40¡. In higher latitudes, fluxes are markedly lower because the sea ice cover artificially prevents them in the model, except where runoff from rivers (northern hemisphere) or ice shelf (southern hemisphere) is prescribed.
III.3.3. Isotopic field and relation to salinity.
The seasonal variability of salinity and oxygen 18 will not be described here since we are interested only in their average characteristics. An important point is that the simulated oxygen 18 - salinity relationship, on a basin scale, appears stable over the seasons. Overall the salinity variability seems underestimated compared to observed one, probably because the coarse resolution of the atmospheric GCM somehow smooths the prescribed fluxes.
Figures III.10c and d show the surface (first model level) salinity and oxygen 18 fields simulated by the GCM after 2200 years. Characteristics similar to those of the atmospheric fluxes (Figures III.10a and b) are obvious : a strong latitudinal gradient results from enriched subtropics and depleted higher latitudes, with intermediate values in the intertropical zone except in the Atlantic which is much more enriched than the two others basins.
Figure III.10. Results from the 2200 year simulation with OPA oceanic GCM and from a 2-box model (where a deep homogeneous reservoir balances the atmospheric fluxes). All fields concern the surface (i.e. the first 10 m level) and are annual averages. Oxygen 18 is expressed as deviation to V-SMOW (in permil). a.
salinity simulated with a box model; b. oxygen 18 simulated with a box model; c.
salinity simulated with the oceanic GCM; d. oxygen 18 simulated with the oceanic GCM; e. salinity compilation from observations (Levitus, 1982).
These features are summarized by the Figure III.11 which compares the latitudinal distribution of the oxygen 18 GEOSECS surface data (1987) with the modelled ones. Although these data are not likely to be representative of the whole ocean, they give an idea of the spatial variability of its isotopic content. There is a quite good agreement between observed and modelled values, especially in the Atlantic although the latter are too enriched there in the subtropics. In the Pacific, the very low modelled values in the intertropical zone correspond to the monsoonal south-eastern Asia characterised by a strong P-E budget (P-E>1.3 m/year). The large number of points there is due to the highest spatial resolution of the model at the lowest latitudes. The Atlantic is clearly more enriched, by more than 0.5 ä, compared to the Pacific.
This difference corresponds to a net freshwater export from the Atlantic to the Pacific basin through Central America, in the real world as well as in our fluxes, since this imbalance is reasonably simulated by the GISS AGCM (1987).
-3 -2 -1 0 1 2
d18 O (permil/ SMOW)
-80 -60 -40 -20 0 20 40 60 80
latitude
OPA
GEOSECS (depth<200m)