Implications of a Carbonate Ion Effect on Shell Carbon and Oxygen Isotopes
519
data for 0. universa. Planktic Ol3C records from
two tropical surface-dwelling species indicate typical interglacial to glacial differences of only -0.2
to -0.3%0 for Globigerinoides sacculifer,
whereas Globigerinoides ruber records indicate
shifts of -0.3%0 up to as large as -0.8%0 (Shackleton
1977; Curry and Crowley 1987; Oppo and
Fairbanks 1989; Mix 1992; Shackleton et al. 1992;
Struck et al. 1993; Linsley and Dunbar 1994; Stott
and Tang 1996). The calculated magnitude of the
carbonate ion effect in warm waters, -0.3 to -0.7%0
(Fig. 2d), falls within the range of observed Ol3C
differences. But in most of the planktic records
from warm waters the minimum ino13C occurs on
the glacial termination, when atmospheric pC0 2
and, by extension, surface water carbonate ion levels were already well on their return to Holocene
values (Sowers and Bender 1995). Another important factor that requires explanation is that the most
positive shell 013C values are often found in the
early Holocene rather than at the core-top. The difference in timing between the tropicaVsub-tropical
013C records and the predicted change in surface
ocean carbonate ion appears to preclude carbonate ion concentrations as the principal control on
shello13C. However, if carbonate ion is an important influence, the paradox of two surface-dwelling tropical species recording such different 013C
shifts could be explained by a greater carbonate ion
response for G. ruber. This hypothesis has recently
been confIrmed by culturing experiments (Bijma et
al. 1998).
Carbonate in the Deep Ocean
Because the deep ocean is not in direct contact
with the atmosphere, there is no single constraint
that can be imposed on the carbonate chemistry of
glacial bottom waters. In addition, it is not known
if carbonate ion concentrations influence the stable isotopic composition of benthic shells. Proposed
scenarios for the glacial abyss range from nearly
unchanged carbonate ion, which would match observations of only small glacial-interglacial changes
in the lysocline (oceanic depth below which sedimentary calcite dissolves), to a near doubling of
deep water carbonate ion concentrations, implied
by proxy pH data (Sanyal et al. 1995). Of the two
hypotheses for reduced glacial pC0 2 considered
here, a strengthened biological pump would lead to
negligible changes in abyssal carbonate ion,
whereas any ofthe CaC0 3 cycling changes would
result in a large increase in abyssal carbonate ion
(Archer and Maier-Reimer 1994).
If there is no carbonate ion influence on benthic
foraminifera, or if deep water carbonate ion remained constant, correcting solely for the effect
of higher glacial carbonate ion on planktic shells
would make planktic 013C values more positive
and therefore increase the true glacial
planktic-benthic 013C difference (il013C). A
greater surface-deep Ol3C difference equates to
a stronger glacial biological pump. A 0.1 %0 increase
in il0 13 C translates to an approximate 10 ppm V
drop in pC0 2 , so a 0.4%0 correction on planktic
Ol3C due to higher carbonate ion combined with the
observed 0.4%0 average glacial increase in il0 13 C
would be suffIcient to explain the entire 80 ppm V
glacial-interglacial pcq difference (Shackleton et
al. 1983a; Shackleton and Pisias 1985; Curry and
Crowley 1987; Shackleton et al. 1992).
The alternative scenario is that carbonate ion
concentration does influence benthic shello 13 C. In
that case, the negative shift in benthico l3 C during
glacial episodes would be consistent with a wholeocean increase in carbonate ion and pH, as expected for a pC0 2 drawdown due to changes in
CaC0 3 cycling (Archer and Maier-Reimer 1994;
Sanyal et al. 1995). Ifthe benthic response is similar to planktics (i.e. --0.01 %0 per flmol/kg), the
-0.4%0 observed change in glacial benthic /i13C
(Curry et al. 1988) would be equivalent to a 40
flmollkg increase in abyssal carbonate ion.
Influence of Carbonate Ion on Glacial
Oxygen Isotope Ratios
Oxygen isotopic ratios in planktic foraminifera are
an important constraint on surface ocean cooling
during the LGM. Isotopic paleotemperatures appear to be consistent with CLIMAP reconstructions, which suggestthat tropical surface temperatures did not cool by more than 2°C during the LGM
(Broecker 1986; Stott and Tang 1996). However,
the extent of agreement depends on the actual
519
data for 0. universa. Planktic Ol3C records from
two tropical surface-dwelling species indicate typical interglacial to glacial differences of only -0.2
to -0.3%0 for Globigerinoides sacculifer,
whereas Globigerinoides ruber records indicate
shifts of -0.3%0 up to as large as -0.8%0 (Shackleton
1977; Curry and Crowley 1987; Oppo and
Fairbanks 1989; Mix 1992; Shackleton et al. 1992;
Struck et al. 1993; Linsley and Dunbar 1994; Stott
and Tang 1996). The calculated magnitude of the
carbonate ion effect in warm waters, -0.3 to -0.7%0
(Fig. 2d), falls within the range of observed Ol3C
differences. But in most of the planktic records
from warm waters the minimum ino13C occurs on
the glacial termination, when atmospheric pC0 2
and, by extension, surface water carbonate ion levels were already well on their return to Holocene
values (Sowers and Bender 1995). Another important factor that requires explanation is that the most
positive shell 013C values are often found in the
early Holocene rather than at the core-top. The difference in timing between the tropicaVsub-tropical
013C records and the predicted change in surface
ocean carbonate ion appears to preclude carbonate ion concentrations as the principal control on
shello13C. However, if carbonate ion is an important influence, the paradox of two surface-dwelling tropical species recording such different 013C
shifts could be explained by a greater carbonate ion
response for G. ruber. This hypothesis has recently
been confIrmed by culturing experiments (Bijma et
al. 1998).
Carbonate in the Deep Ocean
Because the deep ocean is not in direct contact
with the atmosphere, there is no single constraint
that can be imposed on the carbonate chemistry of
glacial bottom waters. In addition, it is not known
if carbonate ion concentrations influence the stable isotopic composition of benthic shells. Proposed
scenarios for the glacial abyss range from nearly
unchanged carbonate ion, which would match observations of only small glacial-interglacial changes
in the lysocline (oceanic depth below which sedimentary calcite dissolves), to a near doubling of
deep water carbonate ion concentrations, implied
by proxy pH data (Sanyal et al. 1995). Of the two
hypotheses for reduced glacial pC0 2 considered
here, a strengthened biological pump would lead to
negligible changes in abyssal carbonate ion,
whereas any ofthe CaC0 3 cycling changes would
result in a large increase in abyssal carbonate ion
(Archer and Maier-Reimer 1994).
If there is no carbonate ion influence on benthic
foraminifera, or if deep water carbonate ion remained constant, correcting solely for the effect
of higher glacial carbonate ion on planktic shells
would make planktic 013C values more positive
and therefore increase the true glacial
planktic-benthic 013C difference (il013C). A
greater surface-deep Ol3C difference equates to
a stronger glacial biological pump. A 0.1 %0 increase
in il0 13 C translates to an approximate 10 ppm V
drop in pC0 2 , so a 0.4%0 correction on planktic
Ol3C due to higher carbonate ion combined with the
observed 0.4%0 average glacial increase in il0 13 C
would be suffIcient to explain the entire 80 ppm V
glacial-interglacial pcq difference (Shackleton et
al. 1983a; Shackleton and Pisias 1985; Curry and
Crowley 1987; Shackleton et al. 1992).
The alternative scenario is that carbonate ion
concentration does influence benthic shello 13 C. In
that case, the negative shift in benthico l3 C during
glacial episodes would be consistent with a wholeocean increase in carbonate ion and pH, as expected for a pC0 2 drawdown due to changes in
CaC0 3 cycling (Archer and Maier-Reimer 1994;
Sanyal et al. 1995). Ifthe benthic response is similar to planktics (i.e. --0.01 %0 per flmol/kg), the
-0.4%0 observed change in glacial benthic /i13C
(Curry et al. 1988) would be equivalent to a 40
flmollkg increase in abyssal carbonate ion.
Influence of Carbonate Ion on Glacial
Oxygen Isotope Ratios
Oxygen isotopic ratios in planktic foraminifera are
an important constraint on surface ocean cooling
during the LGM. Isotopic paleotemperatures appear to be consistent with CLIMAP reconstructions, which suggestthat tropical surface temperatures did not cool by more than 2°C during the LGM
(Broecker 1986; Stott and Tang 1996). However,
the extent of agreement depends on the actual