CPClimate of the PastCPClim. Past1814-9332Copernicus PublicationsGöttingen, Germany10.5194/cp-12-2271-2016The simulated climate of the Last Glacial Maximum and insights into the global marine carbon cycleBuchananPearse J.pearse.buchanan@utas.edu.auhttps://orcid.org/0000-0001-7142-882XMatearRichard J.LentonAndrewPhippsSteven J.https://orcid.org/0000-0001-5657-8782ChaseZannahttps://orcid.org/0000-0001-5060-779XEtheridgeDavid M.https://orcid.org/0000-0001-7970-2002Institute for Marine and Antarctic Studies, University of Tasmania, Hobart, Tasmania, AustraliaCSIRO Oceans and Atmosphere, CSIRO Marine Laboratories, G.P.O. Box 1538, Hobart, Tasmania, AustraliaARC Centre of Excellence in Climate System Science, University of Tasmania, Hobart, AustraliaCSIRO Ocean and Atmosphere, Aspendale, Victoria, AustraliaPearse J. Buchanan (pearse.buchanan@utas.edu.au)22December20161212227122951July201611July201621November201627November2016This work is licensed under a Creative Commons Attribution 3.0 Unported License. To view a copy of this license, visit http://creativecommons.org/licenses/by/3.0/This article is available from https://cp.copernicus.org/articles/12/2271/2016/cp-12-2271-2016.htmlThe full text article is available as a PDF file from https://cp.copernicus.org/articles/12/2271/2016/cp-12-2271-2016.pdf
The ocean's ability to store large quantities of carbon, combined with the
millennial longevity over which this reservoir is overturned, has implicated
the ocean as a key driver of glacial–interglacial climates. However, the
combination of processes that cause an accumulation of carbon within the
ocean during glacial periods is still under debate. Here we present
simulations of the Last Glacial Maximum (LGM) using the CSIRO Mk3L-COAL
(Carbon–Ocean–Atmosphere–Land) earth system model to test the contribution
of physical and biogeochemical processes to ocean carbon storage. For the LGM
simulation, we find a significant global cooling of the surface ocean
(3.2 ∘C) and the expansion of both minimum and maximum sea ice cover
broadly consistent with proxy reconstructions. The glacial ocean stores an
additional 267 Pg C in the deep ocean relative to the pre-industrial (PI)
simulation due to stronger Antarctic Bottom Water formation. However,
889 Pg C is lost from the upper ocean via equilibration with a lower
atmospheric CO2 concentration and a global decrease in export production,
causing a net loss of carbon relative to the PI ocean. The LGM deep ocean
also experiences an oxygenation (> 100 mmol O2 m-3) and
deepening of the calcite saturation horizon (exceeds the ocean bottom) at
odds with proxy reconstructions. With modifications to key biogeochemical
processes, which include an increased export of organic matter due to a
simulated release from iron limitation, a deepening of remineralisation and
decreased inorganic carbon export driven by cooler temperatures, we find that
the carbon content of the glacial ocean can be sufficiently increased
(317 Pg C) to explain the reduction in atmospheric and terrestrial carbon
at the LGM (194 ± 2 and 330 ± 400 Pg C, respectively). Assuming
an LGM–PI difference of 95 ppm pCO2, we find that 55 ppm can be
attributed to the biological pump, 28 ppm to circulation changes and the
remaining 12 ppm to solubility. The biogeochemical modifications also
improve model–proxy agreement in export production, carbonate chemistry and
dissolved oxygen fields. Thus, we find strong evidence that variations in the
oceanic biological pump exert a primary control on the climate.
Introduction
The Late Pleistocene is characterised by a sawtooth-like cycling between cool
glacial and warm interglacial states .
Global temperatures and atmospheric CO2 are strongly correlated across
these climate cycles with approximately 80–100 ppm of change corresponding
to global mean temperature variations of 3–4 ∘C
. This correlation was
extended to causation by evidence that increases in atmospheric CO2
preceded decreases in global ice mass during glacial terminations
. The large potential of the ocean to store
carbon, coupled with evidence that the terrestrial reservoir was diminished
under glacial conditions , has implicated the ocean as
the prime candidate for driving glacial–interglacial changes in atmospheric
CO2 (; ;
). However, identifying the combination of
mechanisms that drove a flux of carbon into the ocean during glacial periods
remains a fundamental and largely unresolved problem.
If we first consider only physical changes, a net influx of CO2 caused by
cooling is a feature of the glacial ocean. However, the increase in
solubility attributed to cooling is partially offset by increased salinity
due to a lower sea level , so that the contribution of
solubility changes is estimated at ∼ 13 ppm of the total 80–100 ppm
CO2 drawdown . Therefore, other physical
changes associated with a glacial climate, notably changes to the large-scale
circulation and sea ice fields , may make a considerable
contribution to carbon sequestration. Proxy evidence and model experiments
of the glacial climate
have shown that a greater proportion of the deep ocean was dominated by
southern source waters seefor a review. The existence
of this glacial-type circulation has been connected to an expanded sea ice
field . An expansion of southern source waters throughout
the deep ocean and an expanded sea ice field are now considered to be primary
candidates for carbon sequestration during glacial climates
.
However, the most promising explanations of the glacial decline in
atmospheric CO2 involve changes to ocean productivity in concert with
reorganisations of the global overturning circulation
. An increased glacial productivity, first
proposed by and explored by by
increasing the global nutrient inventory, is now an established feature of
the subantarctic zone due to enhanced aeolian deposition of iron
. The Southern Ocean exerts a strong control on
atmospheric CO2 through its direct connection with deep waters , and enhanced productivity in
this zone is thus a prime candidate for explaining a fraction of the
glacial–interglacial CO2 difference.
In numerous other regions, however, productivity appears to have been reduced
during glacial climates. The affected regions include waters south of the
Antarctic Polar Front , the North Pacific
, the tropical Indian Ocean
and the equatorial Pacific
. A weaker export production in
these regions would have offset a strengthened biological pump in the
subantarctic, thereby weakening the ability of the ocean to store carbon
during glacial conditions. Whether the strengthening of the biological pump
in the glacial subantarctic was sufficient to increase the carbon content of
the ocean despite losses in productivity in other regions therefore requires
further testing.
This has led some authors to look for alternative biogeochemical mechanisms.
A notable example is the application of temperature-dependent
remineralisation to global fluxes of organics into the interior ocean. The
positive relationship between microbial metabolism and temperature has been
known for some time , but it was only recently that this
relationship was applied to a glacial setting. By prescribing a global
cooling of 5 ∘C, reduced atmospheric CO2 by
∼ 35 ppm. Further research has shown that even a slight deepening in
the remineralisation profile can cause large changes in oceanic carbon
content by reducing surface ocean inorganic carbon concentrations, which in
turn strengthens the air–sea influx of carbon .
Another biogeochemical mechanism that is proposed to increase oceanic carbon
storage is an altered calcium carbonate to organic carbon
(CaCO3 : Corg) export production ratio
. A global decrease in the
CaCO3 : Corg ratio would enhance carbon storage by reducing
the pCO2 of surface waters seefor a review.
A decrease in CaCO3 production could be caused by cooling
and/or an increase in silicate delivery to the lower-latitude oceans that simulated organic carbon production
.
Numerous physical and biogeochemical changes have been associated with a
glacial ocean and all have been identified in some respect as important
drivers of glacial–interglacial climate cycles. Now, recent insights into the
distributions of dissolved oxygen and carbonate species
at the Last Glacial Maximum (LGM; ∼ 21 000 years BCE)
provide new constraints to identify which combination of physical and
biogeochemical changes could have realistically sequestered carbon within the
ocean at this time. Here, we use an earth system model with attached
biogeochemistry, CSIRO Mk3L-COAL (Carbon–Ocean–Atmosphere–Land), to test current theories against these new
insights. Using our simulated LGM ocean state, we quantify the contribution
of physical and biogeochemical changes to the estimated increased of
520 ± 400 Pg of carbon within the oceanic reservoir at the LGM
and demonstrate the importance of marine biogeochemistry
to global climate.
Model and experiments
The model simulations were performed using the CSIRO Mk3L climate system
model version 1.2 , which includes components
that describe the atmosphere, land, sea ice and ocean. The horizontal
resolution of the atmosphere, land and sea ice models is 5.6∘× 3.2∘ in the longitudinal and latitudinal
dimensions, respectively, with 18 vertical levels. The ocean model has a
horizontal resolution of 2.8∘× 1.6∘ with 21
vertical levels. For this study, we conduct simulations using both the full
climate system model and the stand-alone ocean model.
Summary of modelling experiments performed. An O before a model name
denotes that it was an ocean-only simulation. BGC refers to biogeochemistry.
ExperimentModelGreenhouse gasOrbitalCommentforcing (CO2e)aparametersCpl-PICoupled2800 ka BPUnmodified BGCCpl-LGMCoupled16721 ka BPUnmodified BGCExperimentModelAtmosphericClimateCommentCO2 (ppm)stateO-PIOcean280PIUnmodified BGCO-LGMOcean185LGMUnmodified BGCO-PICO2LGMOcean185PIUnmodified BGCO-PIiceLGMOcean185PIUnmodified BGC/LGM sea iceO-PIsolLGMOcean185PIUnmodified BGC/LGM SST and SSSO-LGMpocBGCOcean185LGM10 × potential export productionO-LGMremBGCOcean185LGMIncreased depth of remineralisationbO-LGMpicBGCOcean185LGMNo particulate inorganic carbon exportO-LGMallBGCOcean185LGMBGC modifications of O-LGMpocBGC, O-LGMremBGC and O-LGMpicBGC
a Carbon dioxide equivalents, corresponding to CO2,
CH4 and N2O from 280 ppm, 760 ppb and 270 ppb for PI simulations to
185 ppm, 350 ppb and 200 ppb for LGM simulations. b Power law
exponent for organic matter remineralisation changed from -0.9 to -0.7.
Two fully coupled model experiments were undertaken to simulate the
pre-industrial (Cpl-PI) and Last Glacial Maximum (Cpl-LGM) climates. The
Cpl-PI climate was obtained by forcing the model with an atmospheric CO2
concentration of 280 ppm and by prescribing 1950 CE values for the orbital
parameters. This experiment was integrated for a total of 10 000 years
. The Cpl-LGM simulation followed the protocol developed by
Phase III of the Palaeoclimate Modelling Intercomparison Project (PMIP3),
with the exception that no changes were made to terrestrial topography,
oceanic bathymetry or the positions of the coastlines. The closure of
important oceanic connections due to sea level loss, such as the Bering
Strait, was not considered. The atmospheric CO2 equivalent concentration
was set to 167 ppm, providing a radiative forcing equivalent to the
specified reductions in the atmospheric concentrations of CO2, CH4 and
N2O from 280 ppm, 760 ppb and 270 ppb for pre-industrial simulations to
185 ppm, 350 ppb and 200 ppb for LGM simulations. The orbital parameters
were set to values for 21 000 years BP. The Cpl-LGM simulation was
initialised from the state of the Cpl-PI simulation at the end of model year
100. The model was then integrated for a total of 3900 model years, until it
had reached quasi-equilibrium. Over this integration the ocean experienced an
increase in salinity by 0.5 PSU due to increased evaporation, which
reflected the coupling between a cooler, drier atmosphere and the ocean.
With the climate state provided by the Cpl-LGM experiment, nine additional
ocean biogeochemical simulations were made with different physical and
biogeochemical conditions to explore the effect on the carbon cycle
(Table ). These experiments utilised Mk3L-COAL, an enhanced version
of the Mk3L climate system model which includes biogeochemical modules
embedded within the ocean, atmosphere and terrestrial models. All experiments
were forced by key boundary conditions (wind stresses, temperature, salinity,
incident radiation, sea ice), which were obtained as monthly averages over
the final 50 years of the fully coupled model experiments. The heat and
freshwater fluxes into the ocean were determined by relaxing the sea surface
temperature (SST) and sea surface salinity (SSS) towards the
prescribed fields using a 20-day timescale. An additional 0.5 PSU was added
to the salinity field of the LGM experiments to ensure that the ocean was
1 PSU more saline than the PI. For a description of the ocean
biogeochemistry the reader is directed towards Appendix A of
and the experiments of .
First, five experiments were conducted to explore how the physics of a
glacial ocean affected carbon content. Two experiments, O-PI and O-LGM, were
completed under standard PI and LGM conditions. These standard experiments
were compared with three additional experiments that together were used to
discern how changes in solubility, sea ice and circulation between the PI and
LGM climates affected carbon. One of these experiments,
O-PICO2LGM, involved a standard PI simulation with
prescribed atmospheric pCO2 concentrations of the LGM (185 ppm).
This allowed direct comparison between experiments O-LGM and
O-PICO2LGM to determine how global physical changes
affected carbon storage. To separate the effects of an expanded sea ice
field, solubility changes and circulation changes at the LGM, another two PI
experiments, O-PIiceLGM and
O-PIsolLGM, were completed. These experiments were
forced by PI physics, such that a PI circulation was present, but the
biogeochemical model responded to the sea ice field and sea surface
conditions (temperature and salinity) of the LGM under an atmospheric
pCO2 of 185 ppm. An atmospheric pCO2 of 185 ppm was chosen for direct comparison with
experiment O-PICO2LGM.
Second, four experiments (O-LGMpocBGC,
O-LGMremBGC, O-LGMpicBGC and
O-LGMallBGC) represented LGM experiments in which the
equations controlling biogeochemical cycling were altered. It should be made
clear that these experiments did not explicitly simulate the biogeochemical
changes caused by an altered climate in any mechanistic sense. However, the
prescription of the following changes allowed us to undertake a theoretical
investigation into their capacity to sequester carbon at the LGM. The
experiments were as follows:
O-LGMpocBGC. The scaling factor
(SnppO) was increased by a factor of 10 (Eq. 1) to increase
the export of particulate organic carbon (POC) from the surface ocean and
therefore strengthen the biological carbon pump. Increasing POC export in the
LGM ocean was motivated by an enhanced delivery of iron to the surface ocean
via aeolian dust at the LGM .
POC=SnppO⋅Vmax⋅min[PO4][PO4]+Pk,F(I),
where
Vmax is the temperature-dependent maximum growth of phytoplankton
; Pk is the half-saturation constant for
nutrient-limited growth, set to 0.1 mmol PO4 m-3; and F(I) is
the productivity versus irradiance equation for determining light-limited
growth .
O-LGMremBGC. The POC remineralisation depth
was increased by changing the power law exponent (b) in Eq. (2) from -0.9
to -0.7, which replicated a bulk shift of POC from the upper to the deep
ocean. The motivation for increasing the amount of POC that reaches deeper
levels is the expectation that a cooler ocean would reduce the rate of
bacterial remineralisation in
the upper ocean . This change increased the
simulated POC that reaches the 1000 m depth level from 12.5 to 20 %.
Remin(z)=min(1,(z100)b)
O-LGMpicBGC. Export production of particulate
inorganic carbon (PIC) was turned off by setting the rain ratio
(RPIC) of PIC : POC to zero in Eq. (3). The motivation for
reducing PIC export in the glacial ocean is related to the positive linear
relationship between calcification and temperature and an
increased silicate supply to lower latitudes that potentially favoured
non-calcifying producers .
PIC=POC⋅RPIC
O-LGMallBGC. All three modifications to ocean
biogeochemistry were employed. All ocean-only simulations were integrated for
10 000 years to ensure that the ocean carbon cycle reached a steady state.
To assess whether the behaviour of the biogeochemical tracers within the
coupled model differed from those in the ocean-only model, we ran the coupled
model with online ocean biogeochemistry for a further 1000 years using the
steady-state biogeochemical fields from the ocean-only experiments. This
assessment was made using both the PI and LGM climates. For key diagnostics,
such as the meridional overturning circulation, ocean carbon content and
global export production, the behaviour of the ocean-only simulation differed
by less than 1 % from the coupled simulations. Given the computational
speed of the ocean-only model, these experiments provide an ideal platform to
test the sensitivity of the ocean biogeochemical fields to the
parameterisations used in the biogeochemical model.
Results and discussion
In the following, we first discuss the simulated physical changes to the
ocean observed between the Cpl-PI and Cpl-LGM simulations. Second, we discuss
how the ocean biogeochemical fields differed between the O-PI and O-LGM
simulations, which were forced with the output of the coupled simulations.
Finally, we explore how modifying biogeochemical parameterisations alters the
biogeochemistry, including changes to carbon storage, export production,
carbonate chemistry and dissolved oxygen.
Changes in sea surface temperature (SST), sea ice extent and large-scale circulation between the LGM and PI simulations. SST changes are
compared with proxy reconstructions of SST generated by
and , who use different proxies for
their reconstructions to produce the differences depicted here and discussed
in the text. The values of AABW and NADW formation provided by the models
from the Palaeoclimate Modelling Intercomparison Project Phase II (PMIP2) are
presented by . The transport rate of the Antarctic
Circumpolar Current (ACC) for the PMIP2 models is presented by
. The estimates of NADW production provided by the
PMIP3 models were taken from .
a The rate of Antarctic Bottom Water (AABW) formation
is calculated as the annual average of the most negative rate of overturning
(Sv) in the Southern Ocean south of 60∘ S and deeper than 500 m.
b The rate of overturning in the Atlantic Meridional Overturning
Circulation (AMOC) is calculated as the annual average of the most positive
rate of overturning (Sv) in the North Atlantic Ocean north of 0∘
and deeper than 500 m. c The transport of water by the Antarctic
Circumpolar Current (ACC) is calculated as the annual and zonal average of
the barotropic stream function at 60∘ S.
LGM climate: physical fieldsSea surface temperature (SST)
The simulated change in SST between the Cpl-PI and the Cpl-LGM simulations
shows a similar magnitude and spatial structure to proxy reconstructions and
prior modelling studies, with the greatest cooling in the equatorial oceans, high
latitudes and eastern boundary currents and the least cooling in the
subtropics and western boundary current regions (Fig. ;
Table ). The global SST mean of the Cpl-LGM was 3.2 ∘C
cooler than the Cpl-PI. This change falls within the range of estimates
(∼ 2–4 ∘C) produced by other climate models
but sits towards the cooler limits
of previous multiproxy SST reconstructions that estimate a change of
2 ± 1.8 ∘C . However, a
recent reanalysis of the proxy data presented by
showed past estimates may have underestimated cooling by as much as 50 %
. This finding reconciles some disagreement between climate
models and palaeoproxies and places our simulated cooling of 3.2 ∘C
well within the bounds of uncertainty in reconstructions.
Annual sea surface temperature (SST) difference between
(a) the coupled PI experiment, Cpl-PI and the observations from
and (b) the difference between the coupled LGM
and PI experiments (Cpl-LGM-Cpl-PI). Solid contour lines denote positive
changes in SST by 4 and 8 ∘C, while negative changes in SST are
denoted by dashed lines at 4 and 8 ∘C.
Regionally, the greatest cooling took place in the high latitudes and in the
equatorial Pacific, where temperatures were in excess of 4 ∘C cooler
than the Cpl-PI climate. Meanwhile, the Western Pacific Warm Pool,
subtropical gyres and western boundary currents cooled less
(0.5–3.0 ∘C). Again, proxy and climate
modelling are consistent with both the magnitude and
spatial pattern of cooling. A notable example of model–data agreement is in
the Pacific sector of the Southern Ocean, where SSTs were up to and in excess
of 4 ∘C cooler at the LGM . Enhanced cooling in the
high latitudes and in the eastern boundary currents generated strong zonal
and meridional temperature gradients relative to Cpl-PI SST. There is a
consistent regional pattern to SST cooling in the LGM emerging from proxy and
model simulations that is broadly consistent
with our simulated cooling.
Where there is still large uncertainty in SST change at the LGM is in the
tropical ocean seefor a review. The Cpl-LGM cooling of
3.3 ∘C across the tropical ocean (15∘ S–15∘ N) is
greater than other simulations but falls
well within the -5.1 to -2.17 ∘C estimated by .
Regionally, climate models and proxies
agree that cooling in the tropical Atlantic Ocean
probably exceeded cooling in the tropical Pacific and Indian oceans by
roughly 1 ∘C. In contrast, the tropical Pacific Ocean cooled by
2 ∘C more than the tropical Atlantic and Indian oceans in the
Cpl-LGM simulation. Although SSTs in the east equatorial Pacific have been
reported as 1.5–3.0 ∘C cooler than the PI
, the simulated cooling over much of the
tropical Pacific appears excessive compared to previous simulations
.
Sea ice extent
The Cpl-PI sea ice extent is consistent with estimates made using satellite
measurements during the 1979–1987 period (;
Table ). These measurements represent the first
global estimates of sea ice coverage, and although there is evidence that sea
ice has declined by 20 % since the 1950s , the strong
agreement between the Cpl-PI sea ice fields and the observations of
provide a benchmark for assessing LGM sea ice changes.
Associated with cooler SSTs, sea ice coverage (fractional sea ice area ≥
15 %) was greatly expanded in the Cpl-LGM for both hemispheres relative
to the Cpl-PI (Fig. , Table ). In the Southern Hemisphere,
total sea ice coverage increased by ∼ 120 and ∼ 225 % at its
seasonal maximum and minimum, respectively, relative to the Cpl-PI. In the
Northern Hemisphere, total sea ice coverage increased by ∼ 145 %
and ∼ 105 % at its seasonal maximum and minimum, respectively,
relative to the Cpl-PI. These increases correspond to equatorward expansions
of the sea ice field of between 5–10∘ around the Southern Ocean and
in excess of 15∘ in both the North Atlantic and Pacific oceans.
Annual average sea ice cover for (a) the Cpl-PI Northern
Hemisphere, (b) the Cpl-LGM Northern Hemisphere, (c) the
Cpl-PI Southern Hemisphere and (d) the Cpl-LGM Southern Hemisphere.
The red and blue contour lines in each projection represent the maximum and
minimum seasonal sea ice extents (where sea ice concentration equals 15 %
as per ). In (b), the dashed orange contour
line represents the maximum seasonal sea ice extent produced by the Institut Pierre Simon
Laplace (IPSL) climate system model, which took part in the PMIP3 LGM experiment, and is
broadly consistent with the results of other PMIP3 models. In (d),
the coloured markers represent locations were winter sea ice was deemed to
have been present (blue) and absent (red) at the LGM according to
.
The simulated expansion of sea ice around much of the Southern Ocean agrees
well with proxy reconstructions. Maximum sea ice extent reached as far north
as 47∘ S in both the Atlantic and Indian sectors
and as far north as 55∘ S in the Pacific sector
of the Southern Ocean (;
Fig. ). The magnitude of growth in the Atlantic and Indian sectors has been tested
and largely supported by subsequent studies and
is consistent with our Cpl-LGM sea ice field. In the Pacific sector, however,
the simulated maximum sea ice edge extends well equatorward of the
55∘ S boundary that has been defined by
(Fig. ). By comparing the coverage of sea ice in the Southern
Hemisphere of the Cpl-LGM (∼ 46 × 106 km2) with that
estimated by (∼ 39 × 106 km2), we
can attribute the simulated excess of sea ice in the glacial Southern Ocean
to a possible overestimate in the Pacific sector.
The Cpl-LGM sea ice was broadly consistent with reconstructions in the North
Atlantic, with the exception that too much ice covered the Nordic Seas. The
central and eastern parts of the subpolar North Atlantic, including the
Nordic Seas, are thought to have been at least seasonally ice-free
. The Cpl-LGM sea ice field showed strong,
year-round cover in these regions. However, better model–proxy agreement was produced in other parts of the North
Atlantic. Perennial sea ice cover was present in the Greenland Sea and Fram
Strait during the LGM . There is also
evidence that winter sea ice reached south of Iceland to fill much of the
Labrador Sea and extended along the eastern Canadian
margin . These features were produced in the Cpl-LGM
simulation.
In the North Pacific, the Cpl-LGM sea ice field expanded across a large area.
Proxy reconstructions indicate the presence of strong cover in the Okhotsk
Sea , the Japan Sea and
the western Bering Sea during the LGM, with seasonally
ice-free conditions in the central west . Thus, like the
North Atlantic, sea ice presence was likely more extensive in the western
margins of the North Pacific, and this pattern is replicated by other climate
models (Fig. ). In the Cpl-LGM simulation, more intense sea ice
cover developed in the west, but year-round cover also developed over the
central North Pacific and contrasts directly with the findings of
, who argued for ice-free conditions during the summer.
Thus, our simulated sea ice extent in the North Pacific may be an
exaggeration.
The upper panels depict the total meridional overturning
stream function (Sv) for the global ocean in the (a) Cpl-PI and
(b) Cpl-LGM simulations. The bottom panels depict the total
meridional overturning stream function (sv) for the Atlantic ocean in
the (c) Cpl-PI and (d) Cpl-LGM simulations. Note that those
latitudes corresponding to the Southern Ocean are obscured for (c, d) in the Atlantic Ocean, as these overturning velocities are invalid
considering that waters can exit to the east and west and that the
stream function does not account for these losses.
Meridional overturning circulation
The changes observed in the surface ocean within the Cpl-LGM climate were
accompanied by changes in the global meridional overturning circulation
(Fig. ; Table ). The rate of Antarctic Bottom Water (AABW)
formation in the Southern Ocean doubled between the Cpl-PI and Cpl-LGM
experiments, increasing from 10.3 to 20.2 Sv. An increase in surface
density of 0.9 kg m-3 between 60 to 40∘ S drove this
intensification and also strengthened transport by the Antarctic Circumpolar
Current (ACC) from 140 to 309 Sv. Meanwhile, the Atlantic Meridional
Overturning Circulation (AMOC) weakened from 15.6 to 11.4 Sv. The weakened
glacial AMOC was also associated with a shoaling of its lower boundary from
approximately 3000 to 1500 m. As a result, much of the Atlantic Ocean below
1500 m was dominated by AABW as part of the lower overturning cell.
These changes in the lower and upper overturning cells were conducive to the
development of a global overturning circulation dominated by a denser AABW
and a shallower AMOC. These results are supported by numerous palaeonutrient
tracers showing an increased presence of southern source waters within the
deep North Atlantic Ocean during glacial periods
.
The prevailing interpretation of the palaeonutrient tracers is that the
maximum depth of the AMOC was displaced above about 2000 m and that
this shoaling facilitated the development of a saltier, more stratified
glacial deep ocean . have linked these
changed to the expansion of sea ice in the Southern Ocean, which caused a
greater proportion of Circumpolar Deep Water to rise into a zone of negative
buoyancy flux and thereby produce greater quantities of denser AABW.
However, contradictory changes in the glacial overturning circulation have
been simulated in other climate system models. The rates of AABW formation
tend to increase under LGM conditions , but
responses of the AMOC among models are highly variable. In earlier
experiments as part of the PMIP2 project, the AMOC response ranged between
40 % above and below the PI rate of overturning . Our
weakened (∼ 35 %) and shallower (∼ 50 %) glacial AMOC is
therefore consistent with the lower bounds of these simulations, as are our
rates of AABW formation (Table ). More recent LGM simulations as part
of the PMIP3 project, however, developed stronger and deeper glacial AMOCs
. Furthermore, a recent reconstruction of Southern Ocean
circulation indicates that extreme intensifications of AABW formation and ACC
transport at the LGM are unlikely . Consequently,
these results challenge our simulated changes in meridional overturning, as
well as the prevailing interpretation of palaeonutrient evidence.
Despite inconsistencies between climate model simulations, palaeonutrient
reconstructions continue to support the existence of a shallower AMOC
overlying southern source waters during glacial periods. Variations in the
isotopic signature of Neodynium, for instance, indicate that AABW was more
dominant in the deep ocean during the LGM and that its mixing with a glacial
form of North Atlantic Deep Water (NADW) was more intense . These and other authors
find further support for the presence of a shallower AMOC
above 2500 m. Moreover, simulated distributions of carbon isotopes across a
range of idealised circulations have shown that a shallower AMOC is necessary
to optimise model–proxy agreement
at the LGM . Importantly, our Cpl-LGM simulation developed
an increased presence of AABW throughout the global deep ocean and the
development of a shallower AMOC.
Changes in the concentration of dissolved inorganic carbon
(mmol m-3) of the ocean due to physical differences between the
pre-industrial (PI) and Last Glacial Maximum (LGM). Panels on the left
represent the zonally averaged differences between experiments
(a) O-PICO2LGM and (b) O-LGM with
the O-PI experiment and therefore show total changes in carbon content due
to all physical changes associated with the LGM climate. Panels on the right
represent the depth-averaged differences between experiments
(c) O-PIsolLGM and
(d) O-PIiceLGM with the
O-PICO2LGM experiment and therefore represent the
contribution of solubility and sea ice changes at the LGM to carbon storage
in the ocean.
Global ocean averaged diagnostics from the model simulations
described in Table . The subscript organic refers to the inventory
due to the remineralisation computed from the apparent oxygen utilisation.
POC and PIC refer to the annual export of particulate organic and inorganic
carbon from the upper 50 m, respectively. The tracer columns refer to global
ocean inventory or global ocean mean values. Global inventory of phosphate
was 2.68 Pmol in all simulations.
ModelAtmosphericaPOC PICΔ CarbonbCorgOxygen (mean) Depthc (m)CO2 (ppm)(Pg C year-1) (Pg C year-1)(Pg C) (Pg C)(mmol m-3) where Ωca=1GlobalSouthern OceandGlobalGlobalUppereDeepfGlobalUppereDeepfO-PI2808.021.600.640g0g0g16501811821802666O-PICO2LGM1858.021.600.64-1290-692-59816501811821803234O-PIsolLGM1858.021.600.64-941-435-34613612061952173222O-PIiceLGM1857.961.440.64-1450-759-69116761791811773233O-LGM1854.480.760.36-622-889+2676652802623003208O-LGMpocBGC1855.964.060.47-443-981+5377722712702723297O-LGMremBGC1853.250.750.26-455-790+3357202752652873254O-LGMpicBGC1854.480.760.00-293-451+1586652802623003196O-LGMallBGC1854.863.430.00+317-419+73710042512642372839
a Atmospheric CO2 is prescribed in each experiment.
b Change in the ocean inventory of carbon relative to experiment
O-PI with atmospheric pCO2 at 280 ppm. c Where the
calcite saturation horizon (Ωca=1) exceeds the depth of
the ocean, the lower depth bound of the deepest grid box was included in the
averaging process. d All grid boxes south of 45∘ S.
e All grid boxes above 2000 m depth. f All grid boxes
below 2000 m depth. g Steady-state carbon content of the global,
upper and deep ocean under pre-industrial conditions is 34 114, 17 458 and
16 656 Pg C, respectively.
LGM climate: biogeochemical fields
The physical changes in the ocean between the Cpl-PI and the Cpl-LGM, as
described above, caused significant changes in ocean biogeochemistry within
the ocean-only simulations (Table ). To assist in the discussion of
the large-scale biogeochemical changes, we divide the upper and the deep ocean
based on the 2000 m depth. This approach also allows for more clearly
distinguishing between changes to the global overturning circulation, air–sea
exchange and biological processes on the biogeochemical fields.
Carbon
For experiment O-LGM, the dissolved inorganic carbon (hereafter referred to as
carbon) content of the ocean was 622 Pg C less than in the O-PI experiment
(Fig. ; Table ). The net loss of carbon reflected the
combined effect of physical changes to the ocean, which include an increase
in solubility due to cooling, an expanded sea ice field, an altered
overturning circulation and the tendency for outgassing caused by a lower
pCO2. The physical changes were sufficient in combination to
increase carbon in the deep ocean by 267 Pg C. However, lowering the
atmospheric pCO2 to 185 ppm drove a large amount of carbon out
of the ocean, causing a loss of 889 Pg C from waters in the upper 2000 m.
The combined effect of our simulated LGM physical state could therefore not
overcome the equilibration
with a lower atmospheric pCO2 concentration.
The increase in carbon content in the glacial deep ocean did suggest that
despite the net loss caused by outgassing, the glacial ocean was indeed
conducive to storing carbon. The loss of carbon from the ocean of experiment
O-PICO2LGM demonstrated this (Fig. ;
Table ). The ocean carbon content of
O-PICO2LGM evolved to be 1290 Pg less than in the O-PI
experiment, which placed the O-LGM carbon content as 668 Pg greater than
O-PICO2LGM. This confirmed that the glacial ocean had
a greater ability to store carbon than the PI ocean.
A summary of the total changes in ocean carbon storage and their
drivers between our simulated pre-industrial (PI) and Last Glacial Maximum
(LGM) climates. We use the increase in carbon of our
O-LGMallBGC experiment of 317 Pg C, which falls
within the estimate range of ocean carbon storage at the LGM
, and our simulated loss of 1290 Pg C from the PI ocean
with an atmospheric pCO2 of 185 ppm to determine the
contribution of physical and biogeochemical drivers in ppm of CO2. For
example, an addition of 1607 Pg C in excess of experiment
O-PICO2LGM (+317 Pg C in column 2) would
constitute a contribution of 95 ppm. BGC refers to biogeochemistry.
a Estimate of increase in ocean carbon content during
the LGM made by , whereby atmospheric carbon was reduced by
194 ± 2 Pg C and terrestrial carbon was reduced by
330 ± 400 Pg C. b Assumes all three biological
modifications that were postulated (see Table , experiments
O-LGMpocBGC, O-LGMremBGC and
O-LGMpicBGC) occurred to provide an upper bound
estimate of ocean carbon storage.
To investigate this further, the individual contributions of glacial
solubility, sea ice and circulation to carbon sequestration were determined
by comparing the idealised experiments of O-PIiceLGM
and O-PIsolLGM to O-PICO2LGM.
Under a PI circulation and a pCO2 of 185 ppm, solubility changes
associated with the LGM increased the total carbon content of the ocean by
349 Pg C. Meanwhile, sea ice expansion reduced carbon content by 160 Pg.
The gains and losses of carbon in each experiment occurred in
the high latitudes (Fig. ). Carbon increased markedly in the Arctic
and North Atlantic Ocean due to cooling, while decreases in export production
caused by sea ice expansion reduced carbon content of the subarctic Pacific,
Labrador Sea and across the Southern Ocean. Circulation changes associated
with the glacial climate redistributed carbon from the upper to the deep
ocean and were therefore responsible for the increase in deep-ocean carbon in
O-LGM. The isolated effect of the glacial circulation was the addition of 479
Pg C, which was calculated by solving for the difference between the combined
effect of solubility and sea ice (189 Pg C) and the gain between the O-LGM
and O-PICO2LGM experiments (668 Pg C).
Therefore, we primarily implicate circulation and secondarily solubility
changes as the physical drivers of oceanic carbon storage at the LGM, while
presenting sea ice expansion as a mechanism for reducing ocean carbon
storage. Regarding solubility, previous work has constrained the effect of
glacial solubility on atmospheric carbon to approximately 15 ppm
. By taking the 1290 Pg C difference between
the O-PI and O-PICO2LGM experiments and assuming an a
priori addition of 520 Pg C in a glacial ocean , we can
attribute approximately 18 ppm to our simulated solubility changes
(Table ). Our results are therefore roughly consistent with other
estimates. The
tendency for carbon to be released to the atmosphere would have been further
reduced by storing more carbon in
the deep ocean. We therefore note that circulation changes and cooling would
have had a complementary effect on carbon sequestration and magnified their
individual effects.
Regarding sea ice, a prevailing view is that sea ice expansion would enhance
oceanic carbon storage by limiting air–sea gas exchange .
This theory largely focuses on the restriction of outgassing from carbon-rich
deep waters that upwell in the Southern Ocean. However, this neglects
responses in the Northern Hemisphere and in the biological pump. In a seminal
paper, showed that export production and circulation in
the Antarctic zone have a strong effect on atmospheric CO2. We found that
light limitation of highly productive regions in both hemispheres caused by
increased sea ice cover led to a global loss of 160 Pg C, equivalent to
8 ppm pCO2. Similar responses have been simulated in other
models that consider the impact of sea ice on biological production and do
not consider temperature and salinity changes associated with sea ice growth
(; ).
Our results demonstrated that although the storage of carbon was enhanced in
the glacial ocean (668 Pg C) due to physical changes, namely due to the
overturning circulation, they could not overcome the loss of carbon
(1290 Pg C) caused by equilibration with a lower atmospheric
pCO2.
Nutrients and export production
Like carbon, phosphate (PO4) concentrations in experiment O-LGM were
redistributed from the upper to the deep ocean (Fig. ). This change
was driven by a strengthened AABW formation and is consistent with proxy
reconstructions. Cadmium and δ13C measurements from the Atlantic
Ocean show increased nutrient concentrations in the deep ocean but reduced
levels above 2000 m at the LGM
(; ; ).
Changes in the export production of particulate organic matter (POC)
and Phosphate concentrations between the O-LGM and O-PI experiments.
(a) Annually averaged export of POC from the upper 50 m
(g C m-2 year-1) for O-PI, (b) the
O-LGM-O-PI difference in Phosphate concentrations (mmol m-3) and
(c) the O-LGM-O-PI difference in export production of POC from
the upper 50 m (g C m-2 year-1).
A direct consequence of the redistribution of PO4 was the reduction in the
production of particulate organic matter across many regions of the O-LGM
ocean (Fig. ). With the exception of the South Pacific and isolated
areas in the subtropics, export production in the O-LGM experiment decreased
relative to the O-PI experiment, so that global export production was
56 % of O-PI. The global reduction was also illustrated by a decrease in
regenerated carbon (Corg), which indicated that the biological
carbon pump was weakened (Table ). The reduction in export production
and regenerated carbon for the O-LGM experiment is significant when compared
with other studies that argue for a more efficient glacial biological pump
than that of the Holocene . The weakened
efficiency of our simulated biological pump can be attributed, in part, to a
large decrease in export production from the subantarctic zone. This feature
is in direct conflict with palaeoproductivity proxies in the Atlantic and
Indian sectors of the subantarctic Ocean
and
some parts of the Pacific sector . Outside of
the Southern Ocean, the reduction in export production in the O-LGM
experiment is largely consistent with palaeoproductivity evidence (see
Introduction).
Carbonate chemistry
The loss of phosphate from the upper ocean and its increase at depth was
mirrored by changes in alkalinity and salinity, so that a more alkaline and
saline signature of AABW, relative to NADW, dominated the deep ocean in
experiment O-LGM. Alkalinity and salinity decreased by 66 mmol Eq m-3
and 0.37 PSU in the surface ocean and increased by 147 mmol Eq m-3
and 2.71 PSU in the deep ocean (Fig. ).
Change in the zonally averaged global distribution of
(a) alkalinity (mmol Eq m-3) and (b) salinity (PSU)
between the O-LGM and O-PI experiments (O-LGM-O-PI). Despite the strong
reduction in salinity in the upper ocean of the O-LGM experiment relative to
O-PI, the whole-ocean salt content was greater by 1.0 PSU.
Because the majority of LGM–PI change in salinity occurred in the deep ocean,
the changes in carbonate chemistry across the surface ocean were small.
Little change between experiments O-PI and O-LGM was found in the aragonite
saturation state (Ωar), which is a unitless index indicating
under- and super-saturation at values below and above 1 (Fig. ).
Surface Ωar between 40∘ S and 40∘ N in O-LGM
(Ωar=3.8) was slightly lower than that of O-PI
(Ωar=4.0) but increased in the high-latitude oceans.
Consequently, the simulated Ωar=3.25 isoline, the value at
present used to define the location of viable coral reef conditions
, was nearly unchanged between the O-LGM and O-PI
experiments. Recent sonar and coring in the southern portion of the Great
Barrier Reef have detected the presence of
drowned coral reefs that existed at the LGM as far south as reefs present
today. Such observations are consistent with our O-LGM experiment and
indicates that the extent of viable coral reefs was unlikely to have been
significantly different at the LGM relative to today.
The annual average surface aragonite saturation state
(Ωar) calculated from (a) the observations of
, (b) the O-PI experiment and (c) the O-LGM
experiment.
The depth at which calcite becomes undersaturated in the water
column (where Ωca=1) calculated from (a) the
observations of (global mean calcite saturation
horizon: 2610 m), (b) the O-PI experiment (global mean calcite
saturation horizon: 2666 m) and (c) the O-LGM experiment
(global mean calcite saturation horizon: 3208 m). The contour lines
represent 500, 1000, 2000 and 3000 m depth. Note that the O-LGM experiment,
which is unmodified in its biogeochemistry relative to the O-PI experiment,
is completely saturated for calcite across the majority of the ocean (white
space).
However, the magnitude of increase in alkalinity in the glacial deep ocean,
which was not accompanied by a stoichiometrically matched increase in carbon,
caused unrealistic increases in the calcite saturation horizon
(Ωca=1). The average position of the calcite saturation
horizon increased from 2666 m in O-PI to 3208 m in O-LGM. While this
increase may seem modest, the increase exceeded the ocean floor over the
majority of the ocean, so that seawater was completely saturated for calcite
outside of the eastern tropical Pacific (Fig. ). There is
strong evidence that the carbonate
chemistry of the LGM ocean was not appreciably different to the late Holocene
, and this information places our simulated changes in
deep-ocean Ωca at the LGM as unrealistic.
Dissolved oxygen
Experiment O-PI produced a global average oxygen concentration of
∼ 181 mmol O2 m-3, similar to the PI global average of about
178 mmol O2 m-3. The combination of cooler
SSTs, an enhanced subduction of AABW and the reduction in export production
in experiment O-LGM dramatically increased the oxygen content in both the
upper and deep ocean by ∼ 80 and ∼ 120 mmol m-3,
respectively, which constituted a global increase of 55 %
(Fig. ; Table ).
The increase in dissolved oxygen in O-LGM was considerable but agreed well
with proxy reconstructions for the upper ocean. The oxygen-poor intermediate
waters of the western North Pacific ,
eastern North Pacific
,
eastern South Pacific ,
equatorial Pacific and Indian Ocean
were better oxygenated at
the LGM relative to the PI climate. An important consequence of oxygenating
the upper ocean is a reduction in the strength of denitrification in these regions. Sedimentary δ15N records suggest that global
aggregate rates of water column denitrification rates over the past
200 000 years were lower during glacial periods and higher during
interglacial periods , and this is consistent with the
simulated oxygenation of the upper ocean.
Zonally averaged dissolved oxygen concentrations (mmol m-3) in
(a) the modern ocean according to the World Ocean Atlas
, (b) the O-PI experiment and (c) the
O-LGM experiment.
However, dissolved oxygen concentrations in the deep ocean increased to an
average of 300 mmol O2 m-3 in O-LGM, and this contrasts starkly
with existing palaeoclimate reconstructions. Deep waters of the Indian Ocean , North Atlantic
, Southern Ocean and
equatorial Pacific were poorly ventilated at the LGM
relative to the Holocene. Drawing on a global compilation of similar studies,
and demonstrated that the deep ocean
was largely deoxygenated relative to the Holocene on a global scale. While
the increase in oxygen concentrations in the upper ocean aligned with the
direction of change inferred from proxies, the response in the deep ocean can
be considered unrealistic.
Importance of ocean biogeochemistry for climate
The carbon content, export production field, carbonate chemistry and deep-ocean oxygen content of experiment O-LGM are outstanding in their
disagreement with proxy evidence. Notably, 622 Pg C was lost from
experiment O-LGM relative to O-PI. The standard LGM simulation was therefore
unable to explain the glacial–interglacial drawdown of atmospheric CO2,
despite the existence of a physical ocean state within realistic bounds. If we
are to reconcile the biogeochemistry of the glacial ocean with that inferred
from proxy evidence, we must therefore consider altering ocean
biogeochemistry.
Reconciling the carbon budget
Three plausible modifications to ocean biogeochemistry (see methods) were
considered: (1) increased POC export production, (2) increased depth of POC
remineralisation and (3) reduced PIC export. In the following we step
through the changes to carbon content caused by each modification, and the
reader is directed to Table 3 for reference.
Experiment O-LGMpocBGC. Although the scaling factor
controlling the export production of organic matter was increased 10-fold,
the actual increase in POC export production averaged over the global ocean
was more modest at roughly 30 %. Because most of the ocean became
phosphate limited as greater quantities of nutrients were redistributed into
the deep ocean, the increase in export production in experiment
O-LGMpocBGC was only felt in those regions where PO4
was not limiting. The subantarctic zone of the Southern Ocean experienced the
greatest increase in export production (∼ 250 %), followed by a few
small regions along the Chilean margin and in the northwest Pacific
(Fig. ). These regional responses caused the global net export
production rate to increase from 4.48 to 5.96 Pg C year-1. Although
this rate of POC export production was still lower than the O-PI experiment of
8.02 Pg C year-1, this increased carbon content by 179 Pg C.
Experiment O-LGMremBGC. The shift of organic matter
to depth was associated with a global reduction in POC export production of
∼ 1.2 Pg C year-1 as remineralisation released PO4 and
carbon further from the photic zone. Despite the reduction in the biological
pump, the bulk transfer of POC to depth generated an increase in ocean carbon
storage of 167 Pg C.
Experiment O-LGMpicBGC. The elimination of PIC in
the simulated glacial ocean increased the solubility of CO2 in the surface
ocean and enabled the ocean to store an additional 329 Pg C.
Independently, none of the above modifications were able to increase ocean
carbon content relative to the O-PI experiment
(O-LGMpocBGC: -443 Pg C;
O-LGMremBGC: -455 Pg C;
O-LGMpicBGC: -293 Pg C). However, by employing all
three biogeochemical modifications in one experiment (experiment
O-LGMallBGC), the glacial ocean was able to store an
additional 939 Pg C more than experiment O-LGM and 317 Pg C more than
O-PI.
Change in annually averaged export of particulate organic carbon
from the upper 50 m (g C m-2 year-1) between the LGM and
PI from the experiments with modified biogeochemical formulations for
(a) O-LGMpocBGC- O-PI,
(b) O-LGMremBGC- O-PI and
(c) O-LGMallBGC- O-PI. It should be noted
that the export production field of particulate organic carbon for experiment
O-LGMpicBGC, whereby particulate inorganic carbon was
set to zero, did not differ from unmodified experiment O-LGM and is therefore
not shown. For this comparison, the reader is directed to Fig. 5.
This magnitude of increase places our glacial ocean state within the
plausible bounds required to offset the loss of atmospheric and terrestrial
carbon reported by of ∼ 520 ± 400 Pg C at the
LGM (Table ). By assuming a glacial–interglacial difference in
atmospheric CO2 of 95 ppm and applying this to our changes in carbon
content, we attribute roughly 40 ppm to changes in ocean physics and 55 ppm
to changes in the biological pump. Within the physical changes, 28 ppm is
attributed to the reorganisation of the global overturning circulation, while
12 ppm can be attributed to changes in surface properties, including sea ice
expansion, cooling and salinification.
The depth at which calcite becomes undersaturated in the water
column (where Ωca=1) for the experiments with modified
biogeochemical formulations. (a) O-LGMpocBGC,
(b) O-LGMremBGC,
(c) O-LGMpicBGC and
(d) O-LGMallBGC. The white areas in the ocean
are regions where calcite is completely saturated throughout the water
column.
Reconciling export production
Of the three biogeochemical modifications applied to the LGM ocean, only two
had an effect on POC export, as the amount of PIC exported from the photic
zone has no influence on the amount of POC export. Deepening the
remineralisation of POC (O-LGMremBGC) shifted a greater
fraction of regenerated PO4 into the deep ocean, which resulted in a
global reduction of export production. Increasing the scaling factor
(O-LGMpocBGC), however, caused an increase in global
export production from 4.48 to 5.96 Pg C year-1. Most of this
increase occurred in the Southern Ocean, particularly the subantarctic zone,
and in a few isolated pockets in the northwest Pacific and North Atlantic
where excess nutrients were available (Fig. ).
The increase in the scaling factor dominated the change in export production
produced when combining all three biogeochemical modifications
(O-LGMallBGC). The strong increase in export production
observed in the subantarctic was clearly replicated within this experiment
and reconciles our simulated export production field with current evidence of
productivity at the LGM. In the Southern Ocean, the Atlantic and Indian
sectors of the subantarctic zone experienced a greater flux of organics
. Whether
this was also the case for the Pacific sector remains under debate, with some
evidence for increase and some for no change
. Meanwhile, it is widely accepted that waters
south of the Antarctic Polar Front were reduced in their productivity
,
likely due to increased sea ice extent
and stratification
.
Change in oxygen concentration (mmol m-3) at 500 m between
the LGM and PI for the experiments with modified biogeochemical formulations.
(a) O-LGMpocBGC- O-PI,
(b) O-LGMremBGC- O-PI,
(c) O-LGMpicBGC- O-PI and
(d) O-LGMallBGC- O-PI. A depth of 500 m
is representative of the depth at which the greatest extent of low-oxygen
water exists in the simulated PI climate. It should be noted that the oxygen
field for experiment O-LGMpicBGC, whereby particulate
inorganic carbon was set to zero, did not differ from the unmodified glacial
experiment O-LGM and can therefore be used here as a reference to that
simulation.
In experiment O-LGMallBGC, net export production
remained less than the in O-PI experiment by 3.16 Pg C year-1 despite
the application of biogeochemical modifications. The weakened carbon transfer
to the interior ocean was also observed in the regenerated carbon content of
the ocean (Corg), which was 646 Pg less than in the O-PI
experiment (Table ). The net decline in export production observed in
this study was dominated by the decline in tropical and subtropical waters.
Many palaeoproductivity studies located outside of the subantarctic zone have
found weakened productivity at the LGM
.
Additionally, an enhanced utilisation of available nutrients in the
subantarctic zone would reduce the nutrient
content of intermediate waters formed in the Southern Ocean and would thus
reduce the delivery of nutrients to lower latitudes .
This mechanism coupled with cooler temperatures caused reductions in export
production across much of the mid- and lower-latitude oceans in experiment
O-LGMallBGC, which maintains the qualitative agreement
between the simulated export production field and proxy observations.
However, the global weakening of the biological pump in our simulations is
contrary to proxy and model-based evidence for a strengthened biological pump
(> Corg) at the LGM . Hence,
while we present improved spatial
agreement between O-LGMallBGC and palaeoproductivity
proxies at the LGM, which was essential to increasing the carbon content of
the ocean, we note that additional increases in the export production field
may be valid.
Reconciling carbonate chemistry
There is strong evidence that the
carbonate chemistry of the glacial deep ocean was not appreciably different
to the late Holocene . Because much of the ocean was saturated
for calcite in the O-LGM experiment, additional processes were required to
shoal the depth at which calcite becomes unsaturated (Ωca=1)
and thereby reconcile proxy evidence.
One mechanism to reduce deep-ocean Ωca would be to reduce
continental inputs of alkalinity at the LGM. However, the presence of
glaciers, drier atmospheric conditions and the exposure of continental
shelves due to lower sea level would have increased the supply of carbonates
to the ocean , thereby increasing ocean alkalinity
and further deepening the carbonate saturation horizon. This mechanism has
been largely refuted as having a significant effect on the
glacial–interglacial difference in the carbon budget
and can therefore be ignored.
The individual biogeochemical modifications were also insufficient to reduce
Ωca to be consistent with palaeo evidence. However,
combining all three modifications in experiment
O-LGMallBGC reduced Ωca significantly
(Fig. ) and produced a globally averaged calcite saturation
horizon at 2839 m. Remarkably, our simulated PI to LGM changes in
Ωca are also consistent with palaeoproxy reconstructions in
a regional sense. The calcite saturation horizon in the Pacific Ocean was
deeper in experiment O-LGMallBGC relative to O-PI but
was shallower in the Atlantic Ocean and within the Atlantic and Indian
sectors of the subantarctic zone. Similar changes are seen in the proxy
record, with a deepening of less than 1000 m in the North Pacific and
Southern Ocean at the LGM and a shoaling in
the Atlantic Ocean .
Change in oxygen concentration (mmol m-3) at 3500 m between
the LGM and PI for the experiments with modified biogeochemical formulations.
(a) O-LGMpocBGC- O-PI,
(b) O-LGMremBGC- O-PI,
(c) O-LGMpicBGC- O-PI and
(d) O-LGMallBGC- O-PI. A depth of 3500 m
is representative of the deep ocean. It should be noted that the oxygen field
for experiment O-LGMpicBGC, whereby particulate
inorganic carbon was set to zero, did not differ from the unmodified glacial
experiment O-LGM and can therefore be used here as a reference to that
simulation.
An important caveat of this study, which cannot be ignored, is the exclusion
of calcium carbonate (CaCO3) burial within ocean sediments. Because this
process is not included in the model, it is highly likely that the deepening
of the calcite saturation horizon that occurred in the standard LGM
experiment (O-LGM) was too extreme. CaCO3 burial lowers the alkalinity of
the glacial ocean and is therefore a negative feedback mechanism that
modulates changes in carbonate chemistry seefor a
review. If calcite saturation increases through the water
column, as found in O-LGM, the burial of CaCO3 would increase and
subsequently mitigate the rise in whole-ocean alkalinity, which in turn would reduce calcite saturation and the
ability of the ocean to store carbon. By not taking this process into
account, both the deepening of the calcite saturation horizon and the storage
of carbon were overestimated in experiment O-LGM.
However, the same reasoning can be applied to the experiments with
biogeochemical modifications. If the burial of CaCO3 was included in these
experiments, the shoaling of the calcite saturation horizon would have been
somewhat mitigated by decreased CaCO3 burial that increased ocean
alkalinity . Consequently, the shoaling that was observed
in these experiments was likely exaggerated, just as the deepening observed
in experiment O-LGM was exaggerated. Again, this can be applied to changes in
the carbon content of the ocean, as a shallower calcite saturation horizon
would have reduced CaCO3 burial, thereby increasing whole-ocean alkalinity and the ocean's ability to take up carbon. This effect would have been
particularly important for experiment O-LGMpicBGC,
where inorganic carbon export was eliminated. If whole-ocean alkalinity was
able to respond to a decrease in CaCO3 rain, this would have amplified the
carbon sequestration of experiment O-LGMpicBGC.
Therefore, the exclusion of CaCO3 burial in experiment
O-LGMallBGC (O-LGM) caused an exaggerated shoaling
(deepening) of the calcite saturation horizon and an underestimated
(overestimated) carbon content.
Reconciling dissolved oxygen
As discussed previously, the increase in oxygen concentrations of the upper
ocean in experiment O-LGM is consistent with the current assemblage of proxy
evidence. All experiments with modified biogeochemistry, including
O-LGMallBGC, had little effect on the upper-ocean
oxygen concentration (Fig. ; Table ) and therefore did
not compromise the agreement between simulated and proxy oxygen
reconstructions.
Modifying ocean biogeochemistry did, however, have a large effect on the
oxygen concentrations of the deep ocean (Fig. ). Increasing export
production (O-LGMpocBGC) and deepening the
remineralisation depth (O-LGMremBGC) both reduced
oxygen concentrations by 28 and 13 mmol m-3, respectively. The
combination of these modifications (O-LGMallBGC)
amplified their individual effects, so that deep-ocean oxygen was reduced by
63 mmol m-3 relative to O-LGM. The increased sensitivity of deep-ocean
oxygen to the combination of increased export production and a deeper
remineralisation depth was also observed in the increase in the quantity of
regenerated nutrients (Corg) that resulted (Table ). A
greater proportion of regenerated nutrients relative to preformed nutrients
at the LGM has been identified as a key driver of interior ocean
deoxygenation , and this process was captured
in experiment O-LGMallBGC.
While the combination of biogeochemical modifications
(O-LGMallBGC) did reduce deep-ocean oxygen towards the
concentrations of the PI ocean in a number of areas (Fig. ), by no
means were global deep-ocean concentrations (237 mmol m-3) lower than
those of O-PI (180 mmol m-3). The simulated oxygenation of the deep ocean remains in contrast to proxy records. However, this inconsistency may
be resolved by addressing some key differences in physical and biogeochemical
fields that evolved in our simulations. A key physical mechanism to reduce
deep-ocean oxygen concentrations would include equivalent or slower formation
rates of major ocean deep waters as per combined with
an intensified coverage of sea ice in the region of their formation. The
growth of sea ice and the formation rate of AABW were likely too strong in
our LGM simulation (see Sects. and ), and we
therefore suggest that a sluggish circulation is necessary for reducing deep-ocean oxygen. Key biogeochemical mechanisms include a further increase in
global export production, and/or a different spatial pattern of export
production, and/or increasing the injection of organic matter to deep water
via further lengthening the remineralisation profile. The weakened biological
pump in our LGM simulations contrasts with other studies
and indicates that export production may be
underestimated by our simulations. The combination of a more sluggish deep-ocean circulation with an enhanced export of organics would significantly
reduce the oxygen content of the deep ocean, while potentially further
increasing carbon storage.
Conclusions
In this study we have shown that simulated physical changes in the ocean
state during the climate of the Last Glacial Maximum are not sufficient to
explain the 80–100 ppm drawdown of atmospheric CO2. Physical changes
associated with the glacial climate, including an expanded sea ice field,
increased solubility and a reorganisation of the global overturning
circulation, were responsible for roughly 40 ppm of CO2 drawdown. The
effect of circulation on carbon storage was greatest at 28 ppm and was
associated with an expansion of southern source waters throughout the deep
ocean. Thus, various biogeochemical modifications were necessary to fully explain the atmospheric drawdown of CO2 during the glacial period. The biogeochemical modifications explored in this study were
(1) an increase in export production consistent with greater iron
fertilisation, (2) a shift of remineralisation to the deep ocean consistent
with cooler temperatures, and (3) a decrease in the production of particulate
inorganic carbon consistent with cooler temperatures and increased delivery
of silicate to lower latitudes. Together, these modifications increased the
carbon content of the ocean and, combined with physical changes, were able to
account for the loss of carbon from the atmosphere and land at the Last
Glacial Maximum. Furthermore, their addition improved
model–proxy agreement in the fields
of export production, carbonate chemistry and dissolved oxygen. This study
demonstrates that fundamental changes to ocean biogeochemical function are
required to explain glacial–interglacial cycles.
Data availability
The model output produced during the experiments of this study is held by the
Australian National Computational Infrastructure (NCI) data portal and is
available for download at 10.4225/41/5859eeac6b473
().
Acknowledgements
We wish to thank Katsumi Matsumoto and Andreas Schmittner for their reviews
that significantly improved the manuscript. Funding for this work was
provided by the Australian Climate Change Science Program and CSIRO Wealth
from Ocean Flagship. An award under the Merit Allocation Scheme on the NCI
National Facility at the Australian National University ensured that
numerical simulations could be undertaken. This research was also supported
under the Australian Research Council's Special Research Initiative for the
Antarctic Gateway Partnership (Project ID SR140300001). The authors wish to
acknowledge the use of the Ferret program
(http://ferret.pmel.noaa.gov/Ferret/) for the analysis undertaken in
this work. The matplotlib package , Iris and Cartopy
packages (http://scitools.org.uk/), and cmocean package
were all used for producing visualisations. Edited by: A. Winguth Reviewed by: K.
Matsumoto and A. Schmittner
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