Simulation of Oxygen Isotopes in a Global Ocean Model
683
benthic foraminferal 0 18 0 values from northeast
Atlantic core tops at water depths >2 km. Many
empirically derived paleotemperature equations
have been published during the past several decades (Bemis et al. 1998). In particular, Shackleton
(1974) used core top data to calibrate ao I8 0-temperature relationship for the benthic foraminifer
Uvigerina. At low temperatures it yields larger isotopic differences between calcite and watero l8 0 c
- 0180 w than the one by Erez and Luz (1983). In
contrast, Bemis et al. (1998) derived a new
paleotemperature equation that is also based on
cultured planktic foraminifera and which at low
temperatures yields smaller isotopic differences
(cf. their Fig. 5). A smaller isotope fractionation
effect corresponds to a saturation ofo c at a larger
depth and vice versa. However, our conclusions
would not change if we had chosen any of the other
paleotemperature equations. In particular, the striking difference between Fig. Iia and b would not
be eliminated. With our choice of the
paleotemperature equation, these figures are directly comparable to Figs. 4 and 7 ofZahn and Mix
(1991).
Figure 13 shows the 0P-E distribution, which
reveals a rich meridional structure, quite different
from the constant value assumed by Lehman et al.
(1993) and Mikolajewicz (1998).
Our ultimate goal remains a simulation of the
paleocean and an interpretation of the wealth of
isotope data in terms of water mass changes. In
the case of the LGM the proxy data coverage may
be dense enough to generate the surface forcing
fields from a combination of SST reconstructions,
planktic foraminiferal 0 18 0 data and atmospheric
model output (e.g. Herterich et al. this volume). In
any case, the simulated equilibrium fractionation of
the oxygen isotope composition of calcite ° c can
be directly compared with the benthic foraminiferal
0 18 0 data. Thus the simulation of oxygen isotopes
in an ocean model has the potential to investigate
and validate circulation regimes very different from
the present one.
There is of course a problem in the calculation
of paleosalinity from 0c. It requires linear Ow -salinity relationships that are temporally invariant over
long (geological) timescales (see Section 1). Hence
upon going to the past it is desirable to avoid the
restoring boundary condition on salinity. For this
reason, and for paleo studies of more different periods than the LGM, the ocean model must probably be coupled to a prognostic sea-ice model, or
to both a simple atmosphere and a prognostic seaice model.
A promising application beyond the glacial equilibrium state is the deglaciation phase with rapidly
changingo c values. Lehman et al. (1993)(in a twodimensional ocean model) and Mikolajewicz (1998)
(in a three-dimensional ocean model) indeed employed an idealized Ow tracer in order to compare
meltwater-induced changes in the thermohaline
circulation with the geologic record.
Conclusions
In conclusion,
• a surface 0180 w forcing field can be generated such that a global ocean model reproduces the
GEOSECS 0 18 0 w data fairly well, both along cross
sections in the Atlantic and Pacific oceans and in
terms of regional means,
• a successful simulation of benthic foraminiferal 0 18 0, however, requires a correction of
the warm bias ofthe present global ocean model,
possibly by using an isopycnal transport parameterization,
• 0180 w can be used to characterize different
water masses,
• but since 0 18 0 does not resolve the zone of
c
mixing between NADW and AABW, it would be
desirable to employ another tracer such as 013C
that could reveal additional information on the vertical structure of the deep paleocean,
• for paleostudies in general, the global ocean
model must probably be coupled to simple atmosphere and sea-ice models, such that neither SSS
nor surface 0180 w need to be prescribed and the
use of present-day 0 18 0 w-salinity relationships can
be avoided.
Acknowledgments
We thank Dr. G. Edward Birchfield and Dr.
Christopher D. Charles for their help during the
early stages of this work. We are grateful to Dr.
683
benthic foraminferal 0 18 0 values from northeast
Atlantic core tops at water depths >2 km. Many
empirically derived paleotemperature equations
have been published during the past several decades (Bemis et al. 1998). In particular, Shackleton
(1974) used core top data to calibrate ao I8 0-temperature relationship for the benthic foraminifer
Uvigerina. At low temperatures it yields larger isotopic differences between calcite and watero l8 0 c
- 0180 w than the one by Erez and Luz (1983). In
contrast, Bemis et al. (1998) derived a new
paleotemperature equation that is also based on
cultured planktic foraminifera and which at low
temperatures yields smaller isotopic differences
(cf. their Fig. 5). A smaller isotope fractionation
effect corresponds to a saturation ofo c at a larger
depth and vice versa. However, our conclusions
would not change if we had chosen any of the other
paleotemperature equations. In particular, the striking difference between Fig. Iia and b would not
be eliminated. With our choice of the
paleotemperature equation, these figures are directly comparable to Figs. 4 and 7 ofZahn and Mix
(1991).
Figure 13 shows the 0P-E distribution, which
reveals a rich meridional structure, quite different
from the constant value assumed by Lehman et al.
(1993) and Mikolajewicz (1998).
Our ultimate goal remains a simulation of the
paleocean and an interpretation of the wealth of
isotope data in terms of water mass changes. In
the case of the LGM the proxy data coverage may
be dense enough to generate the surface forcing
fields from a combination of SST reconstructions,
planktic foraminiferal 0 18 0 data and atmospheric
model output (e.g. Herterich et al. this volume). In
any case, the simulated equilibrium fractionation of
the oxygen isotope composition of calcite ° c can
be directly compared with the benthic foraminiferal
0 18 0 data. Thus the simulation of oxygen isotopes
in an ocean model has the potential to investigate
and validate circulation regimes very different from
the present one.
There is of course a problem in the calculation
of paleosalinity from 0c. It requires linear Ow -salinity relationships that are temporally invariant over
long (geological) timescales (see Section 1). Hence
upon going to the past it is desirable to avoid the
restoring boundary condition on salinity. For this
reason, and for paleo studies of more different periods than the LGM, the ocean model must probably be coupled to a prognostic sea-ice model, or
to both a simple atmosphere and a prognostic seaice model.
A promising application beyond the glacial equilibrium state is the deglaciation phase with rapidly
changingo c values. Lehman et al. (1993)(in a twodimensional ocean model) and Mikolajewicz (1998)
(in a three-dimensional ocean model) indeed employed an idealized Ow tracer in order to compare
meltwater-induced changes in the thermohaline
circulation with the geologic record.
Conclusions
In conclusion,
• a surface 0180 w forcing field can be generated such that a global ocean model reproduces the
GEOSECS 0 18 0 w data fairly well, both along cross
sections in the Atlantic and Pacific oceans and in
terms of regional means,
• a successful simulation of benthic foraminiferal 0 18 0, however, requires a correction of
the warm bias ofthe present global ocean model,
possibly by using an isopycnal transport parameterization,
• 0180 w can be used to characterize different
water masses,
• but since 0 18 0 does not resolve the zone of
c
mixing between NADW and AABW, it would be
desirable to employ another tracer such as 013C
that could reveal additional information on the vertical structure of the deep paleocean,
• for paleostudies in general, the global ocean
model must probably be coupled to simple atmosphere and sea-ice models, such that neither SSS
nor surface 0180 w need to be prescribed and the
use of present-day 0 18 0 w-salinity relationships can
be avoided.
Acknowledgments
We thank Dr. G. Edward Birchfield and Dr.
Christopher D. Charles for their help during the
early stages of this work. We are grateful to Dr.