682
Paul eta!.
Experiment A shows the effect of a homogenous
surface 8 w distribution in the Southern Ocean on
the surface 8 fluxes and the zonal mean 8 distribution. Th; zero slope of the b w -S relati;nship
indirectly accounts for the sea-ice effect. As a
result, the bottom water 8 w values are relatively
more enriched, which illustrates that b w can be
used to constrain the role of sea-ice in AABW
formation (cf. Section I). The apparantly better fit
of Experiment B to the GEOSECS b w data results
from a partial cancellation of errors in the representation of AABW.
The effect ofthe warm bias onblSO e is depicted
in Fig. 11. In the upper 2 km where this bias is largest the 8 e values are too light by about 0.5 %0.
Below 2 km depth the simulated 8 e values exhibit
a vertical gradient that is seen neither in the
GEOSECS nor in the benthic foraminiferal blsO
data. In order to simulate this data successfully, the
bias towards warm temperatures must be reduced
significantly.
A warm bias and a diffuse thermocline are two
widespread problems in ocean modelling, which can
probably be overcome by using diffusion oftracers along isopycnals rather than the physically unjustifiable horizontal diffusion of tracers
(Danabasoglu and McWilliams 1995). The formulation by Gent and McWilliams (1990) includes an
additional tracer transport by mesoscale eddies,
which also eliminates the effect ofthe Deacon cell
on the zonal mean tracer distributions. Finally, the
ACC transport through Drake Passage can probably take a more realistic value due to the isopycnal
form stress that supports more drag with weaker
flow.
The isopycnal transport parameterization taken
by itse If still leads to the formation of too co Id and
fresh AABW which can possibly be corrected for
by a more realistic heat flux parameterization
(Large et al. 1997) or coupling to an atmospheric
energy balance model and a sea ice ocean model
(Fanning and Weaver 1996; Mikolajewicz 1998).
The equilibrium fractionation of the oxygen isotope composition of calciteb e is a good water mass
tracer above 2 km depth. Below 2 km depth the b,
values attain a nearly constant level in the presentday Atlantic Ocean and do not allow the separation ofNADW and AABW. The near constancy
ofb, below 2 km depth in the present-day Atlantic
Ocean is an important feature in itself that must be
simulated correctly before the wealth of isotope
data can be interpreted in terms of water mass
changes.
In the present-day North Atlantic, water masses
deeper than 1 km have considerably higher salinity and b w than in the global ocean. This salinity and
be excess persisted during glacial times, although
NADW is probaly produced at a lower rate (Boyle
and Keigwin 1987, Duplessy et al. 1988). Zahn and
Mix (1991) find that the LGM b values from deep
water Atlantic sites (between i and 4 km depth)
are on average 0.2 %ohigher than those from
Atlantic Bottom Water sites (below 4 km depth)see the depth distributions of benthic foraminiferal
8 1s O ofZahn and Mix (1991) (their Fig. 6) and of
(Labeyrie et al. 1992) (their Fig. 3-b). A meridional cross section shows that the deep water be
maximum is most pronounced north of looN (Zahn
and Mix 1991, Fig. 7). It can be traced through the
north-east Atlantic up to Rockall Plateau where it
is reflected in high 8 e values between 1.4 and 2.3
km depth (Jung 1996).
The LGM 8, values from Pacific Deep Water
sites are similar to those from Atlantic Bottom
Water sites. Zahn and Mix (1991) conclude that if
the measured gradients could be taken at face
value, and if the slope of the b w -S relationship were
assumed to be larger at the LGM than at present,
then Atlantic bottom waters and Pacific deep waters would be allowed to have a common source,
again in the Southern Ocean.
A clear 8 e signal can only be seen in the coldest
water at the very bottom of the glacial Atlantic
Ocean. If the zone of mixing between NADW and
AABW rises by a few hundred meters, this can
probaly not be detected in the depth distribution of
b . Therefore it would be desirable to incorporate
a'tracer such as b!3C into the ocean model that
could reveal additional information on the vertical
structure of the deep ocean.
In computing the equilibrium fractionation
of bISO, we used the paleotemperature equation
of Erez and Luz (1983), which is calibrated
between 14°C and 30°C with cultured planktic
foraminifera. Zahn and Mix (1991) found that this
paleotemperature equation gives the best fit to the
Paul eta!.
Experiment A shows the effect of a homogenous
surface 8 w distribution in the Southern Ocean on
the surface 8 fluxes and the zonal mean 8 distribution. Th; zero slope of the b w -S relati;nship
indirectly accounts for the sea-ice effect. As a
result, the bottom water 8 w values are relatively
more enriched, which illustrates that b w can be
used to constrain the role of sea-ice in AABW
formation (cf. Section I). The apparantly better fit
of Experiment B to the GEOSECS b w data results
from a partial cancellation of errors in the representation of AABW.
The effect ofthe warm bias onblSO e is depicted
in Fig. 11. In the upper 2 km where this bias is largest the 8 e values are too light by about 0.5 %0.
Below 2 km depth the simulated 8 e values exhibit
a vertical gradient that is seen neither in the
GEOSECS nor in the benthic foraminiferal blsO
data. In order to simulate this data successfully, the
bias towards warm temperatures must be reduced
significantly.
A warm bias and a diffuse thermocline are two
widespread problems in ocean modelling, which can
probably be overcome by using diffusion oftracers along isopycnals rather than the physically unjustifiable horizontal diffusion of tracers
(Danabasoglu and McWilliams 1995). The formulation by Gent and McWilliams (1990) includes an
additional tracer transport by mesoscale eddies,
which also eliminates the effect ofthe Deacon cell
on the zonal mean tracer distributions. Finally, the
ACC transport through Drake Passage can probably take a more realistic value due to the isopycnal
form stress that supports more drag with weaker
flow.
The isopycnal transport parameterization taken
by itse If still leads to the formation of too co Id and
fresh AABW which can possibly be corrected for
by a more realistic heat flux parameterization
(Large et al. 1997) or coupling to an atmospheric
energy balance model and a sea ice ocean model
(Fanning and Weaver 1996; Mikolajewicz 1998).
The equilibrium fractionation of the oxygen isotope composition of calciteb e is a good water mass
tracer above 2 km depth. Below 2 km depth the b,
values attain a nearly constant level in the presentday Atlantic Ocean and do not allow the separation ofNADW and AABW. The near constancy
ofb, below 2 km depth in the present-day Atlantic
Ocean is an important feature in itself that must be
simulated correctly before the wealth of isotope
data can be interpreted in terms of water mass
changes.
In the present-day North Atlantic, water masses
deeper than 1 km have considerably higher salinity and b w than in the global ocean. This salinity and
be excess persisted during glacial times, although
NADW is probaly produced at a lower rate (Boyle
and Keigwin 1987, Duplessy et al. 1988). Zahn and
Mix (1991) find that the LGM b values from deep
water Atlantic sites (between i and 4 km depth)
are on average 0.2 %ohigher than those from
Atlantic Bottom Water sites (below 4 km depth)see the depth distributions of benthic foraminiferal
8 1s O ofZahn and Mix (1991) (their Fig. 6) and of
(Labeyrie et al. 1992) (their Fig. 3-b). A meridional cross section shows that the deep water be
maximum is most pronounced north of looN (Zahn
and Mix 1991, Fig. 7). It can be traced through the
north-east Atlantic up to Rockall Plateau where it
is reflected in high 8 e values between 1.4 and 2.3
km depth (Jung 1996).
The LGM 8, values from Pacific Deep Water
sites are similar to those from Atlantic Bottom
Water sites. Zahn and Mix (1991) conclude that if
the measured gradients could be taken at face
value, and if the slope of the b w -S relationship were
assumed to be larger at the LGM than at present,
then Atlantic bottom waters and Pacific deep waters would be allowed to have a common source,
again in the Southern Ocean.
A clear 8 e signal can only be seen in the coldest
water at the very bottom of the glacial Atlantic
Ocean. If the zone of mixing between NADW and
AABW rises by a few hundred meters, this can
probaly not be detected in the depth distribution of
b . Therefore it would be desirable to incorporate
a'tracer such as b!3C into the ocean model that
could reveal additional information on the vertical
structure of the deep ocean.
In computing the equilibrium fractionation
of bISO, we used the paleotemperature equation
of Erez and Luz (1983), which is calibrated
between 14°C and 30°C with cultured planktic
foraminifera. Zahn and Mix (1991) found that this
paleotemperature equation gives the best fit to the