658
Paul eta!'
Applications in Paleoceanography
In contrast to present-day oceanography, applications in paleoceanography do not rely on the isotope content of seawater, but on that preserved in
benthic and planktic foraminifera: the equilibrium
fractionation ofthe oxygen isotope composition of
calcite 0 18 0 c or simply 0c' For example, Broecker
(1986, 1989) used benthic and planktic foraminifera
to constrain tropical SSTs and to determine the
salinity contrast between the tropical Atlantic and
Pacific Oceans. Duplessy et al. (1991) present a
reconstruction of Last Glacial Maximum (LGM)
sea surface salinity (SSS) derived from planktic
foraminifera. A number of studies address the vertical structure ofthe ice-age ocean (e.g., Birchfield
1987; Zahn and Mix 1991; Labeyrie et al. 1992).
Lehman et al. (1993) (in a two-dimensional ocean
model) and Mikolajewicz (1998) (in a three-dimensional ocean model) employ an idealized Ow tracer
in order to compare meltwater-induced changes in
the thermohaline circulation with the geologic
record.
The calculation of paleosalinity from Ow requires
linear ow-salinity relationships that are temporally
invariant on short (monthly to seasonal) as well as
on long (geological) timescales. Rohling and Bigg
(1998) critically review this basic assumption and
find that it is unwarranted for many regions ofthe
global ocean because oflocal changes in the freshwater budget and sea-ice coverage that are
advected into the ocean interior. This conclusion
is supported by the estimation of surface Ow from
an atmospheric model (Juillet-Leclerc et al. 1997)
and the simulation of surface Ow in a global ocean
model using full flux boundary conditions at the sea
surface (Schmidt 1998).
We feel attracted to the idea of Zahn and Mix
(1991) to interpret the wealth of oxygen isotope data
in terms of water mass changes. To this end, we
also incorporated Ow into an ocean model, but we
prescribe Ow at the surface and employ restoring
boundary conditions, with the advantage that the
ocean model can be run to equilibrium. The depth
distribution of Ow can then be compaired with the
observations. Since Ow and Tare simulated simultaneously, 0c can be determined from a paleotemperature equation and compared with benthic
foraminiferal 0 18 0.
Here we show the results of a simulation ofthe
present-day ocean and discuss to what degree °
and 0c can be used to characterize different wate;
masses. We find that ° can indeed serve to trace
the formation and spre~ding of deep and bottom
waters.
With respect to 0c we face a fundamental problem: The decrease oftemperature with depth acts
against the decrease of Ow with depth. Hence below 2 km depth in the present-day Atlantic Ocean,
the 0c values attain a nearly constant level and do
not allow to separate deep and bottom waters (cf.
Zahn and Mix 1991, Figs. 4 and 5; however, in the
glacial Atlantic Ocean there might be a small offset of 0.2 %0 between deep and bottom waters). A
technical problem that can probably be solved by
more realistic heat flux and mixing
parameterizations is the warm bias of our ocean
model. This warm bias yields ° c values that are too
light by about 0.5 %0 above 2 km depth and exhibit
a false vertical gradient below 2 km depth.
Our ultimate goal is a simulation of the
paleocean, e.g., at the Last Glacial Maximum
(LGM). This probably requires the ocean model to
be coupled to simple atmosphere and sea-ice models such that neither sea surface salinity (SSS) nor
surface Ow need to be prescribed and the use of
present-day ow-salinity relationships can be
avoided.
Ocean Model
Governing Equations
The ocean model we employ is based on the Modular Ocean Model (MOM, version 2) of the Geophysical Fluid Dynamics Laboratory (GFDL)
(Pacanowski 1996). It is governed by the primitive
equations. These equations are derived from the
basic laws of physics, particularly the conservation
equations for momentum, mass and energy. They
are formally written in spherical polar coordinates:
longitude A increasing to the east, latitude<\> increasing to the north and height z increasing upward.
Paul eta!'
Applications in Paleoceanography
In contrast to present-day oceanography, applications in paleoceanography do not rely on the isotope content of seawater, but on that preserved in
benthic and planktic foraminifera: the equilibrium
fractionation ofthe oxygen isotope composition of
calcite 0 18 0 c or simply 0c' For example, Broecker
(1986, 1989) used benthic and planktic foraminifera
to constrain tropical SSTs and to determine the
salinity contrast between the tropical Atlantic and
Pacific Oceans. Duplessy et al. (1991) present a
reconstruction of Last Glacial Maximum (LGM)
sea surface salinity (SSS) derived from planktic
foraminifera. A number of studies address the vertical structure ofthe ice-age ocean (e.g., Birchfield
1987; Zahn and Mix 1991; Labeyrie et al. 1992).
Lehman et al. (1993) (in a two-dimensional ocean
model) and Mikolajewicz (1998) (in a three-dimensional ocean model) employ an idealized Ow tracer
in order to compare meltwater-induced changes in
the thermohaline circulation with the geologic
record.
The calculation of paleosalinity from Ow requires
linear ow-salinity relationships that are temporally
invariant on short (monthly to seasonal) as well as
on long (geological) timescales. Rohling and Bigg
(1998) critically review this basic assumption and
find that it is unwarranted for many regions ofthe
global ocean because oflocal changes in the freshwater budget and sea-ice coverage that are
advected into the ocean interior. This conclusion
is supported by the estimation of surface Ow from
an atmospheric model (Juillet-Leclerc et al. 1997)
and the simulation of surface Ow in a global ocean
model using full flux boundary conditions at the sea
surface (Schmidt 1998).
We feel attracted to the idea of Zahn and Mix
(1991) to interpret the wealth of oxygen isotope data
in terms of water mass changes. To this end, we
also incorporated Ow into an ocean model, but we
prescribe Ow at the surface and employ restoring
boundary conditions, with the advantage that the
ocean model can be run to equilibrium. The depth
distribution of Ow can then be compaired with the
observations. Since Ow and Tare simulated simultaneously, 0c can be determined from a paleotemperature equation and compared with benthic
foraminiferal 0 18 0.
Here we show the results of a simulation ofthe
present-day ocean and discuss to what degree °
and 0c can be used to characterize different wate;
masses. We find that ° can indeed serve to trace
the formation and spre~ding of deep and bottom
waters.
With respect to 0c we face a fundamental problem: The decrease oftemperature with depth acts
against the decrease of Ow with depth. Hence below 2 km depth in the present-day Atlantic Ocean,
the 0c values attain a nearly constant level and do
not allow to separate deep and bottom waters (cf.
Zahn and Mix 1991, Figs. 4 and 5; however, in the
glacial Atlantic Ocean there might be a small offset of 0.2 %0 between deep and bottom waters). A
technical problem that can probably be solved by
more realistic heat flux and mixing
parameterizations is the warm bias of our ocean
model. This warm bias yields ° c values that are too
light by about 0.5 %0 above 2 km depth and exhibit
a false vertical gradient below 2 km depth.
Our ultimate goal is a simulation of the
paleocean, e.g., at the Last Glacial Maximum
(LGM). This probably requires the ocean model to
be coupled to simple atmosphere and sea-ice models such that neither sea surface salinity (SSS) nor
surface Ow need to be prescribed and the use of
present-day ow-salinity relationships can be
avoided.
Ocean Model
Governing Equations
The ocean model we employ is based on the Modular Ocean Model (MOM, version 2) of the Geophysical Fluid Dynamics Laboratory (GFDL)
(Pacanowski 1996). It is governed by the primitive
equations. These equations are derived from the
basic laws of physics, particularly the conservation
equations for momentum, mass and energy. They
are formally written in spherical polar coordinates:
longitude A increasing to the east, latitude<\> increasing to the north and height z increasing upward.