CPDClimate of the Past DiscussionsCPDClim. Past Discuss.1814-9359Copernicus GmbHGöttingen, Germany10.5194/cpd-11-3597-2015Impact of the oceanic geothermal heat flux on a glacial ocean stateBallarottaM.maxime.ballarotta@natgeo.su.seRoquetF.FalahatS.ZhangQ.https://orcid.org/0000-0002-9137-2883MadecG.Department of Physical Geography, Stockholm University, Stockholm, SwedenBolin Centre for Climate Research, Stockholm University, Stockholm, SwedenDepartment of Meteorology, Stockholm University, Stockholm, SwedenDepartment of Environmental Science and Analytical Chemistry, Stockholm University, Stockholm, SwedenSorbonne Universités (UPMC, Université Paris 06)-CNRS-IRD-MNHN, LOCEAN Laboratory, Paris, FranceNational Oceanography Centre, Southampton, Marine Systems Modelling Group, Southampton, UKM. Ballarotta (maxime.ballarotta@natgeo.su.se)10August2015114359736246July201517July2015This 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/preprints/11/3597/2015/cpd-11-3597-2015.htmlThe full text article is available as a PDF file from https://cp.copernicus.org/preprints/11/3597/2015/cpd-11-3597-2015.pdf
The oceanic geothermal heating (OGH) has a significant impact on the
present-day ocean state, but its role during glacial periods, when the
ocean circulation and stratification were different from those of
today, remains poorly known. In the present study, we analyzed the
response of the glacial ocean to OGH, by comparing ocean simulations
of the Last Glacial Maximum (LGM, ∼21ka ago) including
or not geothermal heating. We found that applying the OGH warmed the
Antarctic Bottom Waters (AABW) by ∼0.4∘C and
increased the abyssal circulation by 15 to 30 % north of
30∘ S in the deep Pacific and Atlantic basins. The
geothermally heated deep waters were then advected toward the Southern
Ocean where they upwelled to the surface due to the Ekman
transport. The extra heat transport towards Antarctica acted to reduce
the amount of sea ice contributing to the freshening of the whole AABW
overturning cell. The global amount of salt being conserved, this
bottom freshening induced a salinification of the North Atlantic and
North Pacific surface and intermediate waters, contributing to the
deepening of the North Atlantic Deep Water. This indirect mechanism is
responsible for the largest observed warming, found in the North
Atlantic deep western boundary current between 2000 and 3000 m
(up to 2 ∘C). The characteristic time scale of the
ocean response to the OGH corresponds to an advective time scale
(associated with the overturning of the AABW cell) rather than
a diffusive time scale. The OGH might facilitate the transition from
a glacial to an inter-glacial state but its effect on the deep
stratification seems insufficient to drive alone an abrupt climate
change.
Introduction
The oceanic geothermal heating (OGH) is the heat flux through the
sea floor which is generated by the internal heat content of the
lithosphere. This flux is maximum near the oceanic ridges or
underwater volcanic regions and is minimum (∼50mWm-2) in the abyssal plains see
e.g..
The importance of the OGH as a heat source for the ocean system has
long been controversial. Although the ocean is largely heated and
thermally driven at the surface, several recent studies suggest
that the OGH can also affect the ocean dynamic and heat budget
. By
applying spatially constant or variable heat flux in Ocean General
Circulation Models (OGCMs) forced with the present day climate, it
is shown that the OGH is a significant forcing that can weaken the
stability of the water column, warm the bottom water and strengthen
the thermohaline circulation (∼5Sv, Sv =106m3s-1).
Sparse observations suggest that the high oceanic heat fluxes are
important factors causing regional bottom water changes on
centennial time scale, such as the observed thermal change in the
abyssal subarctic Pacific Ocean and in the deep
Eurasian Basin of the Arctic ocean , or the rapid
deep water renewal associated with spreading centres
. A recent study based on
laboratory experiment supports the strong effect of the OGH on
local scale but minimises its impact on the thermohaline
circulation and the turbulent mixing .
From a paleo-climate perspective, the abrupt release of potential
energy due to the accumulation of OGH in the deep ocean is invoked
for explaining the rapid temperature variations observed in
reconstructions of the last glacial cycles based on ice and
sediment cores . It is also postulated that the
OGH could have a large impact on the glacial overturning
circulation and the deep water properties, such as the deep
CO2 storage. A climate simulation of the Neoproterozoic
Era (∼700 Ma ago), when the Earth was entirely covered by ice
(the so-called Snowball Earth), reveals that the ocean is not
stagnant and that the OGH may be a driver of its dynamic in
decreasing the density of the abyssal waters, enhancing the
convective vertical mixing and homogeneous temperature and salinity
in the water column .
To our knowledge, the impact of the OGH has not been yet
investigated for more recent glacial climate period, such as the
Last Glacial Maximum (LGM, ∼21ka ago), when the
conditions were colder, the atmospheric CO2 concentration
was lower and the ocean stratification and deep
circulation stronger than those found today see
e.g.,. In the present
study, we explored the impact of the OGH on a glacial ocean state
by using the forced LGM configuration of . We
aimed at (1) evaluating the impact of the OGH on the ocean
circulation, in particular the North Atlantic and abyssal
thermohaline circulation, the advective heat transport and the
stratification; and (2) testing whether the OGH could be a trigger
of transition between glacial and interglacial climate. The paper
is organised as follows: the ocean simulations are described in
Sect. 2; the impact of the OGH on the LGM simulated state is
described in Sect. 3; Results are discussed in Sect. 4 and
a conclusion is given in Sect. 5.
Model description
The NEMO-LIM2 model was used to design the
numerical experiments. The model configuration was similar to the
experiment made by in their study on the
impact of the OGH in the present-day climate. NEMO solves the
primitive equations discretised on a curvilinear horizontal mesh
and was based in our study on a 2∘×2∘
Mercator grid (namely the ORCA2 global configuration). Within the
tropics, the meridional resolution is increased up to
0.5∘. The vertical dimension is discretised into 31
unevenly spaced depth levels (10 m at the surface and
500 m in the deep ocean). The LGM bathymetry is derived
from the present-day bathymetry minus 120 m, representative
of the alteration of the sea level due to the freshwater storage in
the continental ice-sheets during the LGM. The vertical eddy
viscosity and diffusivity coefficients required to model the
vertical mixing were computed from the Turbulent Kinetic Energy
(TKE) turbulent closure model . The NEMO model uses the TEOS-10 equation of state
. The parameterisation of the mesoscale
eddy-induced turbulence was established by the
formulation, which associates an eddy-induced velocity to the
isoneutral diffusion. The ocean is coupled to the Louvain-La-Neuve
Ice Model LIM2 which solves
the thermodynamic growth and decay of the sea ice, the sea ice
dynamic and its transport.
NEMO-LIM2 was initialised at rest and with the temperature,
salinity and sea-ice fields averaged over the last
100 years of a 4000 years long LGM experiment
carried out with the MPI-OM coupled model . The
surface boundary conditions are computed using the CORE bulk
formulae and the atmospheric fields from a LGM
quasi-equilibrated climate model experiment
. Note that no restoring sea surface salinity
term was applied but the freshwater budget was constrained to have
an instantaneous zero global mean value. A more exhaustive
presentation of the experimental setup and boundary conditions can
be found in .
For the present study, we designed a reference experiment (REF)
without OGH at the sea floor. In a second experiment (GH),
spatially varying OGH fluxes were applied as the bottom boundary
condition. Following , the OGH is computed
from the age of the bedrock. We assumed that the OGH flux during
the LGM was the same as today, since it is estimated from the age
of the bedrocks expressed in million of years and that the LGM
continental plate arrangement was similar to the modern day
condition. The total energy input from the OGH forcing is 29.9 TW
(TW =1012 W) and the mean value over the ocean is ∼88mWm-2. This OGH forcing modifies the heat content
by changing the temperature trend in the model grid boxes just
above the ocean floor.
For each configuration, NEMO was integrated for ∼14 000 years and the analysis covered the last 100
model-years. At this stage of the integration, the ocean model was
close to equilibrium. The annual mean model drifts both in
temperature and salinity was weak, with <0.006∘Ccentury-1 (<0.003∘Ccentury-1) below 500 m and <0.09∘Ccentury-1 (<0.08∘Ccentury-1) in the upper 500 m for
REF (GH). The model drift in salinity was also weak for both GH and
REF, with <0.004 PSU century-1 below 500 m and <0.012 PSU century-1 in the upper layer.
ResultsImpact of the geothermal heat flux on the stratification
The annual mean temperature drift induced by the OGH as a function
of depth is shown in Fig. a, b averaged over the Atlantic
and the Indo-Pacific basins, respectively. In the GH experiment,
the Atlantic basin is ∼0.4∘C colder above
1500 m and gains heat below 1500 m with a maximum
warming of ∼0.9∘C formed between 2000 and
3000 m. In the Indo-Pacific basin, the upper 1500 m
layer is ∼0.25∘C warmer, whereas the deep
layer is up to 0.4 ∘C warmer. In the Atlantic
basin, the heat accumulation due to the OGH below 1500 m
has a characteristic time scale of ∼1600years and
reaches an asymptotic limit of 0.38 ∘C
(Fig. c). In the Indo-Pacific basin, the characteristic
time scale is ∼1200years and the accumulated heat
reaches an asymptotic limit of 0.31 ∘C. An
equilibrium is reached after ∼10 000 years.
The annual zonal mean temperature and salinity patterns in REF for
the Atlantic and the Indo-Pacific basins are shown in
Fig. a–d. The deep ocean is filled with cold (near
the freezing point of sea-water) and saline waters, which agrees
with paleo-proxy reconstructions and
the simulation by . Relatively fresh and cold
waters are found between 40 and 90∘ N in the North
Atlantic and North Pacific basins, due to the presence of sea ice.
The impact of the OGH on the zonal mean temperature and salinity
patterns is shown in Fig. e–h. The temperature
differences are significant at all depth except in the upper
200 m where the temperature variability is strong and
mainly controlled by the atmospheric state. The North Atlantic
cooling found in Fig. is mainly associated with colder
surface water in the Nordic Seas (up to 0.7 ∘C
colder) and with the intrusion of colder Antarctic Intermediate
Water (AAIW) in the South Atlantic basin (up to
1.3 ∘C colder). The deep temperatures in the
Atlantic Ocean are up to 1.3 ∘C warmer,
particularly between 1500 and 3000 m in the deep western
boundary current (Fig. ), and between 30 and
45∘ N. In the Indo-Pacific basin, the layer below
1500 m is up to 0.4 ∘C warmer and the
surface layer is slightly colder (0.1 ∘C colder) in
the North Pacific basin and ∼0.3∘C warmer in
the South Pacific. The salinity differences are significant at all
depth and the patterns are relatively similar between the Atlantic
and the Indo-Pacific basins: the Antarctic Bottom Water (AABW) is
∼0.1 PSU fresher in GH than in REF whereas the upper layer is
between 0.1 and 0.3 PSU saltier.
These differences in the temperature and salinity patterns modify
the sea-water density (Fig. ). The AABW becomes less
dense (the density decreases by ∼0.2kgm-3 in
the Indo-Pacific, 0.3 kgm-3 in the Atlantic) due to
warming and freshening, whereas the density increases up to
0.3 kgm-3 in the thermocline due to colder and more
saline waters. The stratification is hence increased by ∼3% near 2250 m and is reduced by ∼3% near 3250 m (not shown).
Impact of the geothermal heat flux on the thermohaline circulation
Most paleo-climate studies investigate the thermohaline circulation
in the latitude-depth coordinates. A better description of the
circulation is however found in latitude-density coordinates
. have also shown that the
strength of the glacial overturning strongly depends on the choice
of coordinate system. Therefore, we present hereafter the
meridional overturning circulation (MOC) in latitude-density
coordinates, more precisely σ4 (i.e., referenced to
4000 m), in order to better capture and compare the abyssal
circulation. Because the OGH affects the ocean's density structure,
subtracting the streamfunctions latitude-density coordinates is
meaningless. It is however possible to identify the maximum of the
AABW transport at each latitude, as well as the density where the
waters are formed. The MOC in latitude-depth coordinates is
presented and discussed in the Appendix for the reader who is not
familiar with the MOC in latitude-density coordinates.
The annual mean effective (Eulerian mean + eddy induced velocities)
MOC in latitude-density coordinates is shown in Fig. ,
for REF and GH. In the Southern Ocean, relatively dense waters
(σ4>46.2kgm-3) are formed between 60 and
80∘ S. These waters are then transported almost
adiabatically up to 40∘ N. The OGH intensifies the AABW
cell by 15–30 % (from 20.5 Sv in REF to
23.4 Sv in GH in the Southern Ocean, 3.0 to 3.7 Sv
in the Atlantic basin, 6.3 to 8.4 Sv in the Indo-Pacific
basin) and shifts the maximum overturning towards lighter density
classes (from σ4∼47.8kgm-3 to
σ4∼47.6kgm-3). The transport
associated with the North Atlantic Deep Water (NADW) is ∼4Sv stronger in GH than in REF along the
46.5–46.7 kgm-3 isopycnals and the maximum of the
North Atlantic overturning is 11 % larger in GH
(17.2 Sv) than in REF (15.4 Sv). The MOC in
latitude-density coordinates also shows the denser NADW in GH than
in REF. Associated with it, the volume of the AABW in the Atlantic
basin is eroded by ∼15% in GH than in REF after
6000 model years (Fig. ) when the layer below
1500 m is warmed by ∼0.3∘C. In the
Indo-Pacific basin, the volume of the AABW is slightly larger
(0.3 %).
Impact on the northward heat transport
The annual mean effective (computed from Eulerian mean + eddy
induced velocities) northward heat transport (in PW =1015 W) for the Global Ocean, the Atlantic basin and the
Indo-Pacific basin in REF is shown in Fig. . In the
Indo-Pacific basin, the heat transport reaches maxima of 1 and
2 PW at 14∘ N and 14∘ S, respectively, and is
directed northward in the Northern Hemisphere, southward in the
Southern Hemisphere. In the Atlantic basin, the heat transport is
directed towards the North pole at all latitude and is maximum
near 22∘ N (∼0.9 PW). In the Southern ocean, the
heat transport is less than 0.5 PW and directed towards
Antarctica. The impact of the OGH on the northward heat transport
is statistically significant (based on a t test, p value less than
5 %) in the Atlantic and Southern oceans. The AABW in the
Indo-Pacific basin gains geothermal heat when it spreads
northward. Most of this heat (∼0.03 PW) is exported to the
Southern Ocean surface where it diverges towards Antarctica (∼0.02 PW) and towards the South Atlantic Ocean (∼0.05 PW)
near 50∘ S. The geothermal heat transported towards
Antarctica then participates in the relative freshening of the
Southern Ocean surface water. The geothermal heat in the South
Atlantic is transported northward and reaches a maximum of ∼0.12 PW near 40∘ N, where the maximum mixed layer depths
are found.
Impact on the North Atlantic deep convection
Our results suggest that the impact of the OGH on the glacial ocean
stratification and thermohaline circulation is significant. The OGH
warms the AABW by ∼0.4∘C and increases the
abyssal circulation between 15 and 30 % north of
30∘ S in the deep Pacific and Atlantic basins. The
geothermally heated deep waters are advected by the deep
overturning cell and upwell at the Southern Ocean surface. When
reaching the Southern Ocean surface, these waters diverge near
50∘ S towards Antarctica and towards the North Atlantic
basin, due to the Ekman transport. The transport towards Antarctica
contributes to the freshening of the surface waters. As a result,
the newly formed AABW becomes less saline. Due to the global salt
conservation, the freshening of the AABW is compensated by more
saline surface waters in the North Atlantic and North Pacific
(∼0.2 PSU saltier), favouring the densification and the
deepening of the NADW. The largest warming is hence found in the
North Atlantic deep western boundary current between 2000 and
3000 m due to the deepening of the
thermocline. Consequently, the volume of AABW is reduced by
15 % in the Atlantic basin. We found that this mechanism
is relatively fast and has a characteristic time scale of ∼1500years. It corresponds mainly to an advective time
scale (associated with the dynamic of the AABW) rather than
a diffusive time scale.
Discussion
The ∼0.4∘C warming of the abyssal ocean due
to OGH is similar to the results found in simulation of the
present-day climate . However, the largest temperature difference is
found between 1500 and 3000 m in the Atlantic basin due to
the deepening of the thermocline. This latest result coincides
with the result found by in their present-day
climate simulations where the North Atlantic deep western boundary
current warms between 0.9 and 1.5 ∘C, but it
contrasts from the solution found in the simulations from
and where the largest
warming takes place in North Pacific below 3000 m
depth. The mechanism, explained above, is compatible with the
results found in . Both in our study and in
, and in opposition to
and other studies, the ocean surface salinity is not relaxed
towards a climatology. Therefore, the warming of the abyssal
waters contributes to freshening of the Southern Ocean surface
waters via the advection of heat. The large formation of AABW
contributes to fresher (reduced salinity) abyssal waters. Due to
the closed freshwater budget and no restoring term in the sea
surface salinity in the model, the Southern Ocean freshwater
supply is counter-balanced by the densification of the surface
waters becoming more saline in the North Atlantic and Pacific
Ocean. As a result, the AMOC is reinvigorated by the increased
surface salinity.
We found that the maximum of the AMOC is ∼11–15 %
larger in GH than in REF, which is similar to the value found in
simulation of the the present-day climate
. This value may be considered as relatively
important in light of the estimation made for the future climate
scenarios (an average reduction of 25 % in
or on short time scale, but it is relatively
weak compared to the variation of the AMOC on climate time scale,
such as the 75 % reduction with respect to LGM period
during Heinrich stadial 1 (∼15–18.5 ka ago), the
45 % reduction during the Younger Dryas stadial (∼12ka ago) , or the values found in fresh
water hosing experiments under LGM conditions (>20%
reduction in ). In these experiments, the
AMOC changes are linked with surface processes, such as the
freshwater discharge which have
a stronger and faster impact on the thermohaline circulation than
the processes induced by the OGH.
Similar to and , we
found that the impact of the OGH on the northward heat transport
is weak (∼10%) but non-negligible, particularly in
the Atlantic Ocean and in the polar regions as a result of the
large scale advection of the abyssal heat content. We found that
the alteration of the ocean heat transport induced by the OGH in
the North Atlantic (∼0.1 PW) is ∼3 time larger than
the total energy input provided by OGH (0.03 PW). It seems that
the Southern Ocean Ekman transport prevents the accumulation of
OGH in the abyssal ocean. For a salinity gradient of ∼1 PSU, a temperature gradient of ∼3∘C
would be required to destabilise the water column (see Appendix
B). In the present study, the OGH warms by ∼0.4∘C. Therefore the OGH alone is not sufficient
to destabilise the water column. It seems that OGH facilitates the
transition from a glacial to an inter-glacial state by reducing
the volume of saline abyssal waters by ∼15% and
reinvigorating the North Atlantic overturning by ∼10%, but it is unlikely that the OGH is the only cause
to abrupt climate changes.
Conclusions
In the present study, we investigated the response of the ocean to
the geothermal heat flux during a glacial period, such as the LGM,
when the ocean circulation and stratification were different from
today. We found that the heat flux at the sea floor is
a significant forcing of the deep ocean and the global
thermohaline circulation. The Antarctic Bottom Water participates
in the transport of geothermally heated waters from the
Indo-Pacific to the North Atlantic basin, indirectly favouring the
deep convection in the North Atlantic and contributing to the
deepening of North Atlantic Deep Water.
The deep ocean circulation and the OGH hence may speed up the
transition from glacial to inter-glacial ocean state by reducing
the volume of saline abyssal waters and reinvigorating the North
Atlantic overturning. However, a new steady-state is achieved only
a few thousands year after OGH is applied wherein the deep
stratification, albeit weakened, remains extremely stable due to
the strong salinity gradient. We thus find it unlikely that abrupt
climate changes could be triggered by the action of OGH alone
during the LGM period. However, the OGH should contribute
significantly in the transition between glacial and inter-glacial
ocean states. The OGH has a strong effect on the ventilation of
the abyssal ocean and might modulate the time scale of the
overturning, and in turn, the rate of CO2 release from
the deep ocean to the atmosphere.
Our results rely on forced (i.e. prescribed atmospheric
conditions) ocean simulation of the LGM period. It thus does not
account for possible ocean feedbacks on the
atmosphere. Sensitivity study with fully coupled
ocean-atmosphere-biochemistry simulations would be useful to
assess the impact of the OGH on the global climate system.
Impact of the geothermal heat flux on the thermohaline circulation in latitude-depth coordinates
A In this section, we present the annual mean effective (Eulerian
mean + eddy induced velocities) meridional overturning circulation
(MOC) in latitude-depth coordinates
(Fig. a, b). The structure of the LGM
thermohaline circulation agrees with the recents findings derived
from multiple paleo-proxies . The circulation representative of the
North Atlantic Deep Water (NADW) in the upper 2000 m has
a maximum transport of ∼17Sv at 900 m depth
near 35∘ N. It is slightly stronger and shallower than in
present-day simulations with same NEMO-ORCA2 model
, due to
a larger intrusion of the AABW in the Atlantic basin.
The difference in the MOC between GH and REF is shown in
Fig. c, d for the Atlantic and Indo-Pacific
basins. The impact of the OGH on the thermohaline circulation is
statistically significant (based on a t test, p value less
than 5 %) in the Atlantic basin, in the Southern Ocean,
in the Arctic basin below 1000 m and in the Indo-Pacific
basin below 3000 m. The volume transport in the
downwelling branch and the deep current of the NADW is up to
5.6 Sv larger. It is mainly associated with the deepening
of the NADW in the GH experiment. The maximum of the AMOC is ∼15% larger in GH (20 Sv) than in REF
(17 Sv). In the Southern Ocean, the volume transport is
∼4Sv larger in upwelling branch of the Deacon Cell,
between 34 and 60∘ S. Note that the Deacon Cell is
fictitious and mainly appears in latitude-depth coordinates. The
Southern Ocean overturning circulation is better described in
latitude-density coordinates than in latitude-depth coordinates
, because it removes
the fictitious Deacon Cell. The volume transport is ∼4.1Sv larger in the deep AABW cell between 45 and
25∘ S and near Antarctica. In the North Atlantic and
North Pacific, the volume transport in the AABW is between 1 and
2 Sv larger.
Ratio between the thermal expansion coefficient (α) and the saline contraction coefficient (β)
The ratio between the thermal expansion coefficient (α) and
the saline contraction coefficient (β) is <13PSU∘C-1 in our simulation. It
corresponds to the compensation of the variation of potential
temperature due to changes of salinity
. Hence, for a salinity gradient of ∼1 PSU, a temperature gradient of minimum ∼3∘C would be required to destabilise the water
column by mixing processes. In the present study, we found that
the OGH warms the deep ocean by only
0.4 ∘C. Therefore the OGH alone is not sufficient
to abruptly destabilise the water column.
Acknowledgements
The authors acknowledge the National Supercomputer Centre at
Linköping University (Sweden) for providing the computational
resources to run the model. The simulations have been run on the
Triolith super-computer
(https://www.nsc.liu.se/systems/triolith/). Many thanks to
Laurent Brodeau for installing the NEMO model on the Triolith
platform.
ReferencesAdcroft, A., Scott, J. R., and Marotzke, J.: Impact of geothermal heating on the global ocean circulation, Geophys. Res. Lett., 28, 1735–1738,
doi:10.1029/2000GL012182, 2001.Adkins, J. F.: The role of deep ocean circulation in setting glacial climates, Paleoceanography, 28, 539–561,
doi:10.1002/palo.20046, 2013.Adkins, J. F. and Schrag, D. P.: Reconstructing Last Glacial Maximum bottom water salinities from deep-sea sediment pore fluid profiles, Earth Planet. Sc. Lett., 216, 109–123,
doi:10.1016/S0012-821X(03)00502-8, 2003.Adkins, J. F., McIntyre, K., and Schrag, D. P.: The salinity, temperature, and δ18O of the glacial deep ocean, Science, 298, 1769–1773,
doi:10.1126/science.1076252, 2002.Adkins, J. F., Ingersoll, A., and Pasquero, C.: Rapid climate change and conditional instability of the glacial deep ocean from the thermobaric effect and geothermal heating, Quaternary Sci. Rev., 24, 581–594,
doi:10.1016/j.quascirev.2004.11.005, 2005.Ahn, J. and Brook, E. J.: Atmospheric CO2 and climate on millennial time scales during the last glacial period, Science, 322, 83–85,
doi:10.1126/science.1160832, 2008.Anderson, R. F., Ali, S., Bradtmiller, L. I., Nielsen, S. H. H., Fleisher, M. Q., Anderson, B. E., and Burckle, L. H.: Wind-driven upwelling in the Southern Ocean and the deglacial rise in atmospheric CO2, Science, 323, 1443–1448,
doi:10.1126/science.1167441, 2009.Ashkenazy, Y., Gildor, H., Losch, M., Macdonald, F. A., Schrag, D. P., and Tziperman, E.: Dynamics of a snowball Earth ocean, Nature, 495, 90–93,
doi:10.1038/nature11894, 2013.Ashkenazy, Y., Gildor, H., Losch, M., and Tziperman, E.: Ocean circulation under globally glaciated snowball Earth conditions: steady-state solutions, J. Phys. Oceanogr., 44, 24–43,
doi:10.1175/JPO-D-13-086.1, 2014.Ballarotta, M., Brodeau, L., Brandefelt, J., Lundberg, P., and Döös, K.: Last Glacial Maximum world ocean simulations at eddy-permitting and coarse resolutions: do eddies contribute to a better consistency between models and palaeoproxies?, Clim. Past, 9, 2669–2686,
doi:10.5194/cp-9-2669-2013, 2013a.Ballarotta, M., Drijfhout, S., Kuhlbrodt, T., and Döös, K.: The residual circulation of the Southern Ocean: which spatio-temporal scales are needed?, Ocean Model., 64, 46–55,
doi:10.1016/j.ocemod.2013.01.005, 2013b.Ballarotta, M., Falahat, S., Brodeau, L., and Döös, K.: On the glacial and interglacial thermohaline circulation and the associated transports of heat and freshwater, Ocean Sci., 10, 907–921,
doi:10.5194/os-10-907-2014, 2014.Björk, G. and Winsor, P.: The deep waters of the Eurasian Basin, Arctic Ocean: geothermal heat flow, mixing and renewal, Deep-Sea Res. Pt. I, 53, 1253–1271,
doi:10.1016/j.dsr.2006.05.006, 2006.
Blanke, B. and Delecluse, P.: Low frequency variability of the tropical atlantic ocean simulated by a general circulation model with mixed layer physics, J. Phys. Oceanogr., 23, 1363–1388, 1993.Brandefelt, J. and Otto-Bliesner, B. L.: Equilibration and variability in a Last Glacial Maximum climate simulation with CCSM3, Geophys. Res. Lett., 36, 1–5,
doi:10.1029/2009GL040364, 2009.Brodeau, L., Barnier, B., Treguier, A.-M., Penduff, T., and Gulev, S.: An ERA40-based atmospheric forcing for global ocean circulation models, Ocean Model., 31, 88–104,
doi:10.1016/j.ocemod.2009.10.005, 2010.Curry, W. B.: Glacial water mass geometry and the distribution of δ13C of ΣCO2 in the western Atlantic Ocean, Paleoceanography, 20, 1–13,
doi:10.1029/2004PA001021, 2005.Davies, J. H. and Davies, D. R.: Earth's surface heat flux, Solid Earth, 1, 5–24,
doi:10.5194/se-1-5-2010, 2010.
de Lavergne, C., Madec, G., Le Sommer, J., George Nurser, A. J., and Naveira Garabato, A. C.: On the consumption of Antarctic Bottom Water in the abyssal ocean, J. Phys. Oceanogr., in revision, 2015.Detrick, R., Williams, D., Mudie, J., and Sclater, J.: The Galapagos spreading centre: bottom-water temperatures and the significance of geothermal heating, Geophys. J. Int., 38, 627–637,
doi:10.1111/j.1365-246X.1974.tb05433.x, 1974.
Döös, K.: Semianalytical simulation of the meridional cells in the
Southern Ocean, J. Phys. Oceanogr., 24, 1281–1293, 1994.
Döös, K. and Webb, D.: The deacon cell and the other meridional
cells of the Southern Ocean, J. Phys. Oceanogr., 24, 429–442, 1994.
Duplessy, J. C., Shackleton, N. J., Fairbanks, R. G., Labeyrie, L., Oppo, D., and Kallel, N.: Deepwater source variations during the last climatic cycle and their impact on the global deepwater circulation, Paleoceanography, 3, 343–360, 1988.Emile-Geay, J. and Madec, G.: Geothermal heating, diapycnal mixing and the abyssal circulation, Ocean Sci., 5, 203–217,
doi:10.5194/os-5-203-2009, 2009.Evans, H. K. and Hall, I. R.: Deepwater circulation on Blake Outer Ridge (western North Atlantic) during the Holocene, Younger Dryas, and Last Glacial Maximum, Geochem. Geophy. Geosy., 9, Q03023,
doi:10.1029/2007GC001771, 2008.Ferrari, R., Jansen, M., Adkins, J., Burke, A., Stewart, A., and
Thompson, A.: Antarctic sea ice control on ocean circulation in
present and glacial climates, P. Natl. Acad. Sci. USA, 111, 8753–8758,
doi:10.1073/pnas.1323922111, 2014.Fichefet, T. and Maqueda, M. A. M.: Sensitivity of a global sea ice model to the treatment of ice thermodynamics and dynamics, J. Geophys. Res., 102, 12609–12646,
doi:10.1029/97JC00480, 1997.Gaspar, P., Grégoris, Y., and Lefevre, J.-M.: A simple eddy kinetic energy model for simulations of the oceanic vertical mixing: Tests at station papa and long-term upper ocean study site, J. Geophys. Res., 95, 16179–16193,
doi:10.1029/JC095iC09p16179, 1990.Gent, P. and McWilliams, J.: Isopycnal mixing in Ocean Circulation models, J. Phys. Oceanogr., 20, 150–155,
doi:10.1175/1520-0485(1990)020<0150:IMIOCM>2.0.CO;2, 1990.Gherardi, J. M., Labeyrie, L., Nave, S., Francois, R., Mc-Manus, J. F., and Cortijo, E.,: Glacial-interglacial circulation changes inferred from 231 Pa/230 Th sedimentary record in the North Atlantic region, Paleoceanography, 24, 1–14,
doi:10.1029/2008PA001696, 2009.Goutorbe, B., Poort, J., Lucazeau, F., and Raillard, S.: Global heat flow trends resolved from multiple geological and geophysical proxies, Geophys. J. Int., 187, 1405–1419,
doi:10.1111/j.1365-246X.2011.05228.x, 2011.Hautala, S. L., Johnson, H. P., and Bjorklund, T.: Geothermal heating and the properties of bottom water in Cascadia Basin, Geophys. Res. Lett., 32, L06608,
doi:10.1029/2004GL022342, 2005.Heinrich, H.: Origin and consequences of cyclic ice rafting in the northeast Atlantic Ocean during the past 130,000 years, Quaternary Res., 29, 142–152,
doi:10.1016/0033-5894(88)90057-9, 1988.Hemming, S. R.: Heinrich events: massive late Pleistocene detritus layers of the North Atlantic and their global climate imprint, Rev. Geophys., 42, RG1005,
doi:10.1029/2003RG000128, 2004.Hieronymus, M. and Nycander, J.: The budgets of heat and salinity in NEMO, Ocean Model., 67, 28–38,
doi:10.1016/j.ocemod.2013.03.006, 2012.Hofmann, M. and Maqueda, M.: Geothermal heat flux and its influence on the oceanic abyssal circulation and radiocarbon distribution, Geophys. Res. Lett., 36, L03603,
doi:10.1029/2008GL036078, 2009.Joyce, T. M., Warren, B. A., and Talley, L. D.: The geothermal heating of the abyssal subarctic Pacific Ocean, Deep-Sea Res., 33, 1003–1015,
doi:10.1016/0198-0149(86)90026-9, 1986.Kageyama, M., Merkel, U., Otto-Bliesner, B., Prange, M., Abe-Ouchi, A., Lohmann, G., Ohgaito, R., Roche, D. M., Singarayer, J., Swingedouw, D., and X Zhang: Climatic impacts of fresh water hosing under Last Glacial Maximum conditions: a multi-model study, Clim. Past, 9, 935–953,
doi:10.5194/cp-9-935-2013, 2013.
Large, W. G. and Yeager, S. S.: Diurnal to decadal global forcing for ocean and sea-ice models: the data sets and flux climatologies. NCAR Technical Note,
NCAR/TN-460+STR, Boulder, Colorado, CGD Division of the National Center for Atmospheric Research, 2004.
Lecointre, A.: Variabilité interannuelle à décénale en
atlantique nord et mers nordiques: etudes conjointe d?observations,
simulations numériques et réanalyses, PhD thesis, Université
Joseph Fourier, Grenoble 1, 2009.Lippold, J., Luo, Y., Francois, R., Allen, S. E., Gherardi, J., Pichat, S., Hickey, B., and Schulz, H.: Strength and geometry of the glacial Atlantic Meridional Overturning Circulation, Nat. Geosci., 5, 813–816,
doi:10.1038/ngeo1608, 2012.Lynch-Stieglitz, J., Adkins, J. F., Curry, W. B., Dokken, T., Hall, I., Herguera, J. C., Hirschi, J., Ivanova, E., Kissel, C., Marchal, O., Marchitto, T. M., McCave, I. N., McManus, J. F., Mulitza, S., Ninnemann, U., Peeters, F., Yu, E. F., and Zahn, R.: Atlantic meridional overturning circulation during the Last Glacial Maximum, Science, 316, 66–69,
doi:10.1126/science.1137127, 2007.
Madec, G.: NEMO ocean engine, Note du Pôle de modélisation de l'Institut Pierre-Simon Laplace No. 27, Institut Pierre-Simon Laplace, Paris, France, 2008.Marchitto, T. M. and Broecker, W. S.: Deep water mass geometry in the glacial Atlantic Ocean: a review of constraints from the paleonutrient proxy Cd/Ca, Geochem. Geophy. Geosys., 7, Q12003,
doi:10.1029/2006GC001323, 2006.Mashayek, A., Ferrari, R., Vettoretti, G., and Peltier, W. R.: The role of the geothermal heat flux in driving the abyssal ocean circulation, Geophys. Res. Lett., 40, 3144–3149,
doi:10.1002/grl.50640, 2013.McDougall, T. J.: Neutral density surface in the ocean: implications for modelling, Geophys. Res. Lett., 14, 797–800,
doi:10.1029/GL014i008p00797, 1987.
Meehl, G. A., Stocker, T. F., Collins, W. D., Friedlingstein, P.,
Gaye, A. T., Gregory, J. M., Kitoh, A., Knutti, R., Murphy, J. M.,
Noda, A., Raper, S. C. B., Watterson, I. G., Weaver, A. J., and
Zhao, Z.-C.: Global climate projections, in: Climate Change 2007: The
Physical Science Basis. Contribution of Working Group I to the Fourth
Assessment Report of the Intergovernmental Panel on Climate Change,
edited by: Solomon, S., Qin, D., Manning, M., Chen, Z., Marquis, M., Averyt, K. B., Tignor, M., and Miller, H. L., Cambridge University Press, Cambridge, UK and New York, NY, USA, 2007.Monnin, E., Indermühle, A., Daellenbach, A., Flueckiger, J.,
Stauffer, B., Stocker, T. F., Raynaud, D., and Barnola, J.-M.: Atmospheric CO2 concentrations over the Last Glacial Termination, Science, 291, 112–114,
doi:10.1126/science.291.5501.112, 2001.Otto-Bliesner, B. L., Hewitt, C. D., Marchitto, T. M., Brady, E. C., Abe-Ouchi, A., Crucifix, M., Murakami, S., and Weber, S. L.: Last Glacial Maximum ocean thermohaline circulation: PMIP2 model intercomparisons and data constraints, Geophys. Res. Lett., 34, 1–6,
doi:10.1029/2007GL029475, 2007.Petit, J. R., Jouzel, J., Raynaud, D., Barkov, N. I., Barnola, J.-M., Basile, I., Bender, M., Chappellaz, J., Davis, M., Delaygue, G., Delmotte, M., Kotlyakov, V. M., Legrand, M., Lipenkov, V. Y., Lorius, C., Pepin, L., Ritz, C., Saltzman, E., and Stievenard, M.: Climate and atmospheric history of the past 420,000 years from the Vostok ice core, Antarctica, Nature, 399, 429–436, 1999.
Ritz, S. P., Stocker, T. F., Grimalt, J. O., Menviel, L., and Timmermann, A.: Estimated strength of the Atlantic overturning circulation during the last deglaciation, Nat. Geosci., 6, 208–212, 2013.
Roquet, F., Madec, G., McDougall, T. J., and Barker, P. M.: Accurate polynomial expressions for the density and specific volume of seawater using the TEOS-10 standard, Ocean Model., 90, 29–43, 2015.Sarnthein, M., Winn, K., Jung, S. J. A., Duplessy, J. C., Labeyrie, L., Erlenkeuser, H., and Ganssen, G.: Changes in east Atlantic deep-water circulation over the last 30,000 years – 8 time slice reconstructions, Paleoceanography, 9, 209–267, 1994.Scott, J., Marotzke, J., and Adcroft, A.: Geothermal heating and its influence on the meridional overturning circulation, J. Geophys. Res., 106, 31141–31154,
doi:10.1029/2000JC000532, 2001.Siegenthaler, U., Stocker, T. F., Monnin, E., Lüthi, D., Schwander, J., Stauffer, B., Raynaud, D., Barnola, J.-M., Fischer, H., Masson-Delmotte, V., and Jouzel, J.: Stable carbon cycle-climate relationship during the last Pleistocene, Science, 310, 1313–1317,
doi:10.1126/science.1120130, 2005.Skinner, L. C., Fallon, S., Waelbroeck, C., Michel, E., and Barker, S.: Ventilation of the deep Southern Ocean and deglacial CO2 rise, Science, 328, 1147,
doi:10.1126/science.1183627, 2010.Stein, C. and Stein, S.: A model for the global variation in oceanic depth and heat flow with lithospheric age, Nature, 359, 123–129,
doi:10.1038/359123a0, 1992.Tagliabue, A., Bopp, L., Roche, D. M., Bouttes, N., Dutay, J.-C., Alkama, R., Kageyama, M., Michel, E., and Paillard, D.: Quantifying the roles of ocean circulation and biogeochemistry in governing ocean carbon-13 and atmospheric carbon dioxide at the last glacial maximum, Clim. Past, 5, 695–706,
doi:10.5194/cp-5-695-2009, 2009.Timmermann, R., Goosse, H., Madec, G., Fichefet, T., Ethe, C., and Dulière, V.: On the representation of high latitude processes in the ORCA-LIM global coupled sea ice-ocean model, Ocean Model., 8, 175–201,
doi:10.1016/j.ocemod.2003.12.009, 2005.Toggweiler, J. R., Russell, J. L., and Carson, S. R.: Midlatitude westerlies, atmospheric CO2, and climate change during the ice ages, Paleoceanography, 21, PA2005,
doi:10.1029/2005PA001154, 2006.Urakawa, L. and Hasumi, H.: A remote effect of geothermal heat on the global thermohaline circulation, J. Geophys. Res., 114, C07016,
doi:10.1029/2008JC005192, 2009.Watson, A. J. and Garabato, A. C. N.: The role of Southern Ocean mixing and upwelling in glacial–interglacial atmospheric CO2 change, Tellus B, 58, 73–87,
doi:10.1111/j.1600-0889.2005.00167.x, 2006.
Worthington, L.: Genesis and evolution of water masses, Meteor. Mon., 8, 63–67, 1968.Zhang, X., Lohmann, G., Knorr, G., and Xu, X.: Different ocean states and transient characteristics in Last Glacial Maximum simulations and implications for deglaciation, Clim. Past, 9, 2319–2333,
doi:10.5194/cp-9-2319-2013, 2013.Zhou, S., Qu, L., Zhao, X., and Wan, W.: Laboratory simulation of the influence of geothermal heating on the interior ocean, Acta Oceanol. Sin., 33, 25–31,
doi:10.1007/s13131-014-0512-8, 2014.
Annual mean potential temperature drift (in ∘C) induced by the geothermal heat forcing as a function of depth averaged in (a) the Atlantic basin and (b) the Indo-Pacific basin. Contour intervals are every 0.1 ∘C. Note the vertical scale is increased in the upper 1000 m. (c) Time-series of the mean temperature accumulation (in ∘C) due to the geothermal heat below 1500 m in the Atlantic and Indo-Pacific basins. τ1∼1200years and τ2∼1600years denote the characteristic time scale, i.e. the amount of time required for the response to reach (1-1/e)≈63% of the maximum heat accumulation, in the Indo-Pacific and the Atlantic basins.
Annual zonal mean potential temperature patterns (in
∘C) in the reference experiment (REF) for (a)
the Atlantic basin, (b) the Indo-Pacific basin (Contour
interval every 1 ∘C, thick white contour is
0 ∘C); salinity patterns (in ∘C) in the
reference experiment (REF) for (c) the Atlantic basin,
(d) the Indo-Pacific basin (Contour interval every 0.1 PSU, thick white contour is 36.3 PSU); the temperature difference between REF and GH for (e) the Atlantic basin, and (f) the Indo-Pacific basin (Contour interval every 0.1 ∘C, thick grey contour is 0 ∘C); and salinity difference between REF and GH for (g) the Atlantic basin, and (h) the Indo-Pacific basin (Contour interval every 0.1 PSU, thick white contour is 0 PSU). The thick vertical black line shows the location of the South Atlantic entrance at 34∘ S. Note the vertical scale is increased in the upper 1000 m. The patterns in each Southern Ocean sector are shown in each panel between 80 and 34∘ S. The dashed contours represent the region where the difference is insignificant at a 95 % confidence level (based on a t test). AABW: Antarctic Bottom Water, AAIW: Antarctic Intermediate Water.
Map of the annual mean temperature difference (in ∘C) between GH and REF at 2500 m. Maximum and minimum values are denoted in the lower left corner. The largest warming is in the Atlantic deep western boundary current.
Difference in the annual mean and zonal mean potential density patterns (σ4 in kg m-3) between REF and GH for (a) the Atlantic basin, and (b) the Indo-Pacific basin. Contour and scale same as in Fig.
Annual mean density-binned effective (Eulerian + eddy-induced velocities) meridional overturning circulation (in Sv) in the experiment without geothermal heating (REF) for (a) the Atlantic basin, and (b) the Indo-Pacific basin; and in the experiment with geothermal heating (GH) for (c) the Atlantic basin, and (d) the Indo-Pacific basin. The thick black line shows the location of the South Atlantic entrance at 34∘ S. The annual mean meridional overturning circulation in the Southern Ocean is shown in each panel between 80 and 34∘ S. Positive (negative) contours represent clockwise (anti-clockwise) circulations. Contour interval is every 4 Sv, and -1 and +1Sv contours are added. Density bins intervals every 0.01 kgm-3. Note the vertical scale is increased for σ4>46kgm-3.
Time-series of the AABW volume variation (in %) in the Atlantic and Indo-Pacific basins. The AABW volume is computed as the volume of water below 2000 m where the annual mean effective overturning circulation in latitude-depth coordinates is negative (Fig. a, b).
Annual mean effective (Eulerian mean + eddy-induced velocities) northward heat transport (in PW =1015 W) in the Global Ocean, the Atlantic and the Indo-Pacific basins in the reference experiment (dashed line); and the difference between REF and GH in the annual mean northward heat transport (thick line). Note that the difference is magnified by a factor 10. The green dots show where the difference is significant at a 95 % confidence level (based on a t test).
Annual mean effective (Eulerian mean + eddy-induced velocities) meridional overturning circulation in latitude-depth coordinates (in Sv =106m3s-1) in the reference experiment (REF) for (a) the Atlantic basin, and (b) the Indo-Pacific basin. Contour interval is every 4 Sv, and the 0 Sv contours is added. Positive (negative) values represent clockwise (counter-clockwise) circulation. Difference in the effective meridional overturning circulation between REF and GH for (c) the Atlantic basin, and (d) the Indo-Pacific basin. Contour interval is every 1 Sv. The annual mean meridional overturning circulation and the difference in the meridional overturning circulation in the Southern Ocean between GH and REF is shown 80 and 34∘ S. Note the vertical scale is increased in the upper 1000 m. The dashed contours represent the region where the difference is insignificant at a 95 % confidence level (based on a t test).