Climate Dynamics (2004) 22: 293–306
DOI 10.1007/s00382-003-0378-5
Y. Donnadieu Æ G. Ramstein Æ F. Fluteau Æ D. Roche
A. Ganopolski
The impact of atmospheric and oceanic heat transports
on the sea-ice-albedo instability during the Neoproterozoic
Received: 24 February 2003 / Accepted: 17 October 2003 / Published online: 28 January 2004
Ó Springer-Verlag 2004
Abstract In order to simulate the climatic conditions of
the Neoproterozoic, we have conducted a series of simulations with a coupled ocean–atmosphere model of
intermediate complexity, CLIMBER-2, using a reduced
solar constant of 6% and varied CO2 concentrations. We
have also tested the impact of the breakup of the supercontinent Rodinia that has been hypothesized to play an
important role in the initiation of an ice-covered Earth.
Our results show that for the critical values of 89 and
149 ppm of atmospheric CO2, a snowball Earth occurs in
the supercontinent case and in the dislocated configuration, respectively. The study of the sensitivity of the
meridional oceanic energy transport to reductions in CO2
concentration and to the dislocation of the supercontinent demonstrates that dynamics ocean processes can
modulate the CO2 threshold value, below which a
snowball solution is found, but cannot prevent it. The
collapse of the overturning cells and of the oceanic heat
transport is mainly due to the reduced zonal temperature
gradient once the sea-ice line reaches the 30° latitudinal
band but also to the freshening of the tropical ocean by
sea-ice melt. In term of feedbacks, the meridional
atmospheric heat transport via the Hadley circulation
plays the major role, all along the CO2 decrease, by
increasing the energy brought in the front of the sea-ice
margin but does not appear enough efficient to prevent
the onset of the sea-ice-albedo instability in the case of
the continental configurations tested in this contribution.
Y. Donnadieu (&) Æ G. Ramstein Æ D. Roche
Laboratoire des Sciences du Climat et de l’Environnement,
CE-Saclay, Orme des Merisiers, 91191, Gif sur Yvette, France
E-mail: [email protected]
F. Fluteau
Institut de Physique du Globe de Paris, T24-25 E1,
4 place Jussieu, 75252 Paris cedex 05, France
A. Ganopolski
Potsdam Institute for Climate Impact Research,
PO BOX 60 12 03, 14412 Potsdam, Germany
1 Introduction
The extreme Neoproterozoic conditions provide
expanding areas of opportunities to test and to examine
our knowledge of the complex Earth’s system. Recent
interpretations of the geologic record from 750 to
580 Ma argue for periods with a global glaciated Earth,
i.e., the Snowball Earth hypothesis (Hoffman and Schrag 2002). Indeed, new paleomagnetic evidence from
South Australia and Northwest Canada have confirmed
the low-latitude setting for some of the Neoproterozoic
glacial deposits (equatorward of 10°) (Park 1997; Sohl
et al. 1999). Moreover, these glaciations are accompanied by profound modifications of the biogeochemical
cycles as recorded in the Sr and C isotopic anomalies
(Kaufman et al. 1993, 1997). The snowball Earth
hypothesis also provides plausible explanations for the
unusual occurrence of cap carbonates resting above
successive Neoproterozoic glacial deposits. Global glaciations could also act as an environmental filter, given
the harsh climatic conditions, promoting the rapid
evolution of multicellular organisms during the Cambrian explosion as it is seen in the rock record (Hoffman
et al. 1998).
The possible existence of an Earth with global seaice cover and continental ice cover raises fundamental
questions about the mechanisms that would be
responsible for such an event. Budyko (1969) and
Sellers (1969) using energy balance models (EBMs)
were the first to demonstrate that a reduction of the
solar constant by a few percent would lead to a completely ice-covered Earth. One of the prominent features of these conceptual models is the existence of a
fundamental instability caused by the ice-albedo feedback, i.e., if more than about half the Earth’s surface
area were to become ice-covered, the albedo feedback
would be unstoppable. However, as noted by Bendsten
(2002), the strength of this mechanism depends on the
meridional energy transport in the atmosphere and
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Donnadieu et al.: The impact of atmospheric and oceanic heat transports on the sea-ice-albedo instability
ocean which are lumped together in Budyko-Sellers
type EBM.
Since then, the use of atmospheric general circulation
models (AGCMs) that present the advantage to be
quasi-deductive (also called comprehensive) models became widespread. These models then have been used to
simulate the Neoproterozoic glaciations using different
set of boundary conditions. For example, snowball
conditions are found for boundary conditions of
100 ppm atmospheric CO2, 6% reduction in solar forcing and a rectangular supercontinent centered on the
equator representative of the paleogeography for
the Sturtian glacial period (around 750 Ma), when using
the GENESIS version 1.02 AGCM (Jenkins and Smith
1999). In contrast, other modelling studies have focused
on the younger snowball event (i.e. the VarangianVendean glacial period between 620 and 580 Ma) using
more realistic plate configurations. A global glaciated
Earth is accomplished with the GENESIS version 2.0
AGCM (Hyde et al. 2000) but not with the GISS
AGCM (Chandler and Sohl 2000). However these
experiments were carried out with different initial conditions, for instance Hyde et al. (2000) prescribe and fix
land ice from a coupled EBM/Ice Sheet model run
whereas Chandler and Sohl (2000) start from desertcoverage on land. One of the common point of all these
experiments is that they have been made with a slab
ocean that represents the mixed layer but does not take
into account the dynamical behaviour of the ocean. In
this way, Poulsen et al. (2001) using boundary conditions similar to Jenkins and Frakes (1998), but with a
fully coupled ocean–atmosphere general circulation
model (the FOAM model) demonstrated that ocean
dynamics processes can prevent a snowball solution.
However, the duration of their OAGCM simulations are
only 100 years which is short when compared to the time
response of the ocean. This major limitation comes from
the high computational cost of the OAGCM that describes many details of the flow pattern, such as weather
systems and regional currents in the ocean. Finally, the
studies with AGCMs have separately explored some of
the parameter space to identify the region susceptible to
be at the initiation of a snowball Earth (atmospheric
CO2 level, solar constant, high obliquity or paleogeographic control). However, as the computational cost of
these models limits the number of experiments, none of
them has assessed the atmospheric and oceanic heat
transport responses to a progressive decrease of the
greenhouse gases. Only the conceptual EBMs have been
used in this sense, generally to study the faint young sun
climatic paradox (Gérard et al. 1992). The Earth climate
system Model of Intermediate Complexity (EMIC),
CLIMBER-2 (Ganopolski et al. 1998; Petoukhov et al.
2000), provides the possibility to study the whole
parameter space. The model, CLIMBER-2, describes a
large set of processes and feedbacks, but due to low
spatial resolution and simplified governing equations,
has a fast turnaround time. It integrates an atmospheric
module and an ocean module including a sea-ice
module. This coupled model does explicitly resolve the
energy transport in the ocean and in the atmosphere and
thus, is well designed to focus on the interactions between the atmosphere and the ocean that take place
during the cooling of the Neoproterozoic. Since the
simple hypothesis invoking higher obliquity cannot explain all Neoproterozoic glaciations (Donnadieu et al.
2002), the cooling which occurred during these episodes
may be related to changes in atmospheric greenhouse
gas concentrations. Therefore, the purpose of this study
is to explore the effects of changed radiative forcing due
to CO2 decrease and the two different land configuration
on the atmospheric and oceanic heat transports during
the Neoproterozoic using the CLIMBER-2 model. In
particular, we will address the following questions: (1)
how sensitive is the CO2 threshold value (below which a
global glaciated Earth is achieved) to the paleogeography used? (2) What is the response of the atmospheric
and oceanic meridional heat transports during the all
period of the Neoproterozoic cooling? (3) How are those
responses modified when using another paleogeography?
The climate model and the experimental setup are
described in Sect. 2. After briefly showing the main climatic features of our experiments, we focus our analysis
on the changes in the atmospheric and oceanic heat
transports due to the decrease of the atmospheric CO2
levels and to the location of the continental configuration (Sect. 3). The main conclusions of this work are
discussed and summarized in Sect. 4.
2 The model and the experimental setup
2.1 The coupled climate model of intermediate complexity
CLIMBER-2
The model used in this study has been fully described in Petoukhov
et al. (2000). CLIMBER-2 could be classified as a model of intermediate complexity, placed between simple models, usually 1- or
2-D, on one hand, and 3-D climate GCMs on the other hand
(Claussen et al. 2002). The atmospheric module is a 2.5-dimensional statistical-dynamical model which includes many of the
processes that are also described by more sophisticated GCMs. In
contrast with GCMs, it has a coarse resolution of 10° in latitude
and approximately 51° in longitude (Fig. 1a). The large-scale circulation (e.g., jet streams and Hadley circulation) and the main
high and low pressure area are explicitly resolved. The atmospheric
circulation as well as energy and water fluxes are computed at 10
pressure levels, while longwave radiation is calculated using 16
levels. To apply the model to the Neoproterozoic experiments, this
module has been kept as in its present form. The ocean module is
based on the (averaged) equations of Stocker et al. (1992) and
describes the zonally averaged characteristics for three separate
ocean basins (Atlantic, Pacific and Indian) with a latitudinal resolution of 2.5°. It has 21 levels in the vertical including an upper
mixed layer of 50 m thickness. Temperature, salinity, and vertical
and meridional velocity are calculated within each basin. Where the
basins are connected (Fig. 1), the zonal velocities are prescribed.
The model includes a thermodynamic sea-ice model predicting the
sea-ice fraction and thickness for each grid box, a simple treatment
of horizontal transport is also included. The ocean and atmospheric modules are linked by a coupler which calculates the fluxes
of energy, momentum and water between the atmosphere and the
ocean. Each atmospheric model grid box consists of one or several
of six surface types (open water, sea-ice, trees, grass, bare soil and
Donnadieu et al.: The impact of atmospheric and oceanic heat transports on the sea-ice-albedo instability
Fig. 1 Continental reconstructions with the CLIMBER-2 grid:
a present-day, b the Rodinia supercontinent (SC, supercontinent
configuration) at around 800 Ma (adapted from Weil et al. 1998,
Torsvik et al. (2001)). c Configuration after the Rodinia breakup
(DC, dislocated configuration) at around 750 Ma, with location of
the Proto-Pacific Ocean after Meert (2003). The thick dark lines
show the separations in between the three oceanic basins
glaciers). Soil processes are described within a two-layer soil model.
CLIMBER-2 successfully simulates the main features of the modern climate (Ganopolski et al. 1998; Petoukhov et al. 2000) and
compares well with GCMs results for the quaternary climate
(Kubatzki et al. 2000; Ganopolski et al. 2001). Moreover, the
model sensitivity to a CO2 doubling experiment, 2.6 °C, is in the
middle of the range of the GCM experiments (Ganopolski et al.
2001).
2.2 Experimental setups
During the Neoproterozoic, a supercontinent commonly referred
to as Rodinia, supposedly formed at 1100 Ma and broke apart at
around 800–700 Ma. The geometry of the supercontinent Rodinia
is based on the hypothesis that Laurentia formed the core of a
clustering of most of continents (Hoffman 1991) with East
Gondwana situated along its western margin, and with Amazonia
and Baltica positioned along its eastern margin (Dalziel 1997). The
low to mid-latitudes preponderance of land masses was the first
condition proposed by Kirschvink (1992) to favour a global glaciation as it would result in a higher albedo in the tropics. In
addition, Hoffman et al. (1998) recognized that the fragmentation
of the supercontinent may have contributed to the CO2 draw down
by creating many new continental margins (which favours the
burial of organic carbon), then creating amenable conditions to a
global glaciation. Although the exact geometry of that supercontinent at different time slices remains a hard task given the limitations of paleomagnetic data (Meert and Torsvik 2003), we have
chosen two contrasted paleogeographic maps available for the time
period 800–700 Ma, one that is based on the hypothesis of the
Rodinia supercontinent (SC, Fig. 1b, adapted from Weil et al. 1998
and Torsvik et al. 2001) and the other that is one of the
295
hypothetical fragmented configuration (DC, Fig. 1c, adapted from
Meert 2003; Meert and Torsvik 2003). The choice of these two
reconstructions has been made to see if one of those reconstructions favours the onset of a snowball Earth.
Although our goal is obviously not to use the most realistic
continental configuration because (1) it is not available and (2) our
model spatial resolution is rather poor, we only want to quantify
the impact of these two different geometries, on the climate simulated when using a reduced solar constant and different level of
CO2. The total land area of both reconstructions is almost the
same, 100 millions of km2, which represent 20% of the total area of
the Earth. It is probable that the missing additional land masses
was scattered in small terranes but we have also to keep in mind
that those reconstructions do not take into account the shortening
due to continental drift. The elevation for the continents is 100 m
for both configurations. In addition, because river drainage basins
are unknown for this time period, we defined a river mask in which
rain falling and snow melt on land is equally redistributed to all
coastal land points. The surface type, bare soil, is imposed at every
land grid point of the climate model since land plants had yet to
evolve. Albedos of bare soil in the visible and near-infrared
wavebands are 0.18 and 0.36, respectively. Maximum thicknesses of
snow and sea-ice of 5 m and 10 m respectively are imposed in the
CLIMBER-2 model. In fact, the sea-ice is allowed to build up in
the oceanic surface layer. The ocean model bathymetry is a flatbottom case (5000 m depth), the three basins are interconnected
poleward of 50°N and of 60°S for both configurations (Fig. 1b, c)
where zonal advection and diffusion are represented. We preferred
to keep the three-basins structure of the ocean model because it
allows to reproduce a longitudinal resolution of the World Ocean.
We have chosen the longitudinal distribution of each basin in order
to conserve the influence of the continents on the ocean realm (and
vice versa); for instance, we do have a variability between the
western (Fig. 1b, basin 1) and the eastern coastal areas (Fig. 1b,
basin 2) and we also keep a super-ocean (Fig. 1b, basin 3). The
choice of the basins in the DC was inherent to the paleogeography
itself. As there are absolutely no constraints on the nature of the
circumpolar currents during the Neoproterozoic, we use the parametric law which controls the zonal transport between the Atlantic,
Indian and Pacific basins in the CLIMBER-2 model (Ganopolski
et al. 1998; Petoukhov et al. 2000) for the latitudes where the basins
are interconnected (above 50°N and 60°S for both reconstructions).
This hypothesis seems to us the most reasonable (see below for a
discussion of this limitation and a sensitivity study to the zonal
transports). We then assume existence of continental boundaries
which allows to have a longitudinal pressure gradient from 60°S to
50°N.
Boundary conditions also include a solar luminosity 6% below
present (Gough 1981). Milankovitch cycles of the Earth’s orbital
dynamics have been documented in records for the Pleistocene, but
the periods and magnitude of this forcing before the Pleistocene are
uncertain. Hence, we assume the Earth’s orbit about the Sun was
circular (eccentricity = 0) and that the Earth’s obliquity was 23.5°.
This setting causes an equal receipt of solar insolation for both
hemispheres. The numerical experiments were conducted in the
following way: the model was run until equilibrium (5000 years)
under an atmospheric CO2 concentration of 5000 ppm, then
starting from the previous equilibrium state each time, we conducted a suite of simulations (5000 years for each) in which CO2
levels were decreased until a global glaciation was reached. These
settings allow to simulate the transition from a ‘hot’ climate to a
‘cold’ one.
3 Climate simulation results
3.1 Overview
Results from the experiments discussed are summarized
in Table 1. Globally averaged, annual surface air
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Table 1 Overview of the
simulations
SC denotes the experiments
using the Rodinia supercontinent as paleogeography, and
DC denotes those using the
fragmented configuration after
breakup. The following numbers indicate the atmospheric
CO2 level. These abbreviations
are used throughout the manuscript
Simulations Surface air Global temperature Hemispheric fraction
Hemispheric fraction of
temperature of the ocean (°C)
of sea-ice covering the
sea-ice covering the
(°C)
Northern Hemisphere (%) Southern Hemisphere (%)
CTRL
SC3500
SC500
SC200
SC150
SC90
SC89
DC3500
DC500
DC200
DC150
DC149
13.96
13.5
1.91
–8.36
–10.00
–15.09
–42.12
14.42
2.6
–10.0
–13.45
–42.5
4.76
3.01
1.54
–0.5
–0.71
–1.45
–1.8
3.07
0.91
–1.05
–1.43
–1.8
temperatures in the Neoproterozoic simulations
encompass a broad spectrum of values that range from
14.4 to 15.9 °C (Table 1). For both paleogeographies, a
CO2 concentration of 3500 ppm is required to lead to
global air temperature comparable to modern (13.5 and
14.4 °C for the SC and DC experiments respectively).
This is roughly the value we can expect from a single
radiative calculation when we estimate the required CO2
level to overcome the solar intensity reduction. Each
steady-state simulations (with a fixed CO2 value) are
used to build a curve relying the evolution of the globally averaged, annual surface air temperature against the
atmospheric CO2 levels. In this way, two curves are
plotted in Fig. 2, one for the suite of SC runs and one for
the suite of DC runs. The calculated behaviour is similar
to that predicted by EBM/ice-sheet (Hyde et al. 2000)
and AGCM (Baum and Crowley 2001) studies as the
transition from a cold state to a snowball Earth state is
abrupt. In short, as the CO2 levels decrease, the global
temperature first drops quasi linearly until a curvature
appears for a CO2 level of 1000 ppm in both continental
configurations. This non-linearity accentuates below
1000 ppm until the threshold CO2 values that still yields
an ice free tropical ocean, 90 ppm in the SC case and
150 ppm in the DC case, are reached. Then, a CO2 decrease of 1 ppm results in the sudden extension of the
sea-ice line to the equator and in a temperature decrease
of about 25 °C. Hence, the sea-ice-albedo instability
appears at a higher CO2 level when the continents are
fragmented in the CLIMBER-2 model, 149 ppm versus
89 ppm (the reasons for that will be explored in Sect.
3.4).
We start our analysis by presenting a comparison
between the CTRL (i.e. the present-day climate) and the
SC3500 experiments which have nearly the same global
mean temperatures. The only reason to compare such
different climate is to point out the re-organisation of the
ocean and atmosphere transport in a Neoproterozoic
world before the cooling occurs. Then, in order to assess
the relative role of the atmosphere and the ocean when
the instability takes place in the coupled system, we
examine the sensitivity of the oceanic and atmospheric
heat transports to atmospheric CO2 levels in the SC case.
3.8
12.3
27.4
36.9
38
41.9
100
10.8
26.9
39.7
43.2
100
5.3
8.5
26.3
41.5
43.6
50.2
100
5.4
26.1
45.7
49.6
100
Fig. 2 Mean annual globally averaged surface air temperatures for
various CO2 levels from 5000 ppm to 90 ppm (150 ppm) for the
SC simulations (DC simulations)
Finally, the role of the continental configuration is
investigated by comparing the climatic features of the
SC and DC configurations with the same CO2 level.
3.2 Comparison of the SC3500 climate and the presentday climate
The first striking difference using a Neoproterozoic
paleogeography is the absence of continents at high
latitude (Fig. 1). Figure 3a shows the zonally mean annual surface air temperatures for the SC3500 case and
the CTRL case and the difference between the two climates. The SC3500 climate is more equable than the
modern one in the Southern Hemisphere (SH thereafter)
and is characterised by an asymmetry between both
hemispheres with the Northern Hemisphere (NH thereafter) being the coldest (14.4 °C in the NH compared to
16.7 in the SH). Concerning the smaller equator-to-pole
temperature gradient in the SH: (1) the differences in
temperatures ranges from a few degrees Celsius at midlatitudes to 20 °C at polar southern latitudes and is
correlated with the absence of the Antarctica elevated ice
Donnadieu et al.: The impact of atmospheric and oceanic heat transports on the sea-ice-albedo instability
cap in the SC case. (2) The desert conditions specified at
each land points in the SC case result in higher albedo in
the tropics than at present which in turn exhibit colder
temperatures. In the NH, the SC simulated climate is
colder than the modern one at all latitudes (Fig. 3a). The
larger sea-ice cover (Table 1) and the associated higher
albedo in the SC experiment are responsible for this
difference. The seasonal temperature amplitude (not
shown) is almost the same than in the modern case for
the SH but not in the NH where the amplitude of the
variation is smaller. This can be explained by the large
seasonal variation in snow cover for the present day
climate that does not exist in the SC experiment as there
is no continents at mid to high latitudes. In short, the
Fig. 3 a Mean annual zonally averaged surface air temperatures
(°C) for the CTRL run (black line) and the SC3500 run (black
dashed line) and the difference, i.e. the SC3500 run minus the CTRL
run (grey line). b Comparison between atmospheric annual
meridional heat transport calculated for the CTRL run (black line)
and the SC3500 run (black dashed line) (PW). c Advective (grey
lines) and synoptic (black lines) component of the atmospheric
meridional heat transport for the CTRL run (solid lines) and for the
SC3500 run (dashed lines)
297
two main differences between the CTRL and the SC
experiments are the apparition of a flattened temperature gradient in the SH and the strong asymmetry of the
mean temperature between both hemispheres (14.4 °C in
the NH compared to 16.7 in the SH).
In order to better understand the SC3500 climate, we
have to examine the changes in the meridional oceanic
and atmospheric heat transports (OHT and AHT,
thereafter). The response of the yearly averaged meridional AHT is illustrated in Fig. 3b. The atmospheric
circulation plays a major role in carrying energy from
the equator toward the high latitudes to compensate the
net outgoing flux at polar latitudes (Peixoto and Oort
1992). The Hadley cells carry the energy from low to
subtropical latitudes, this transport corresponds to the
advective component in the CLIMBER-2 model
(Fig. 3c). At mid-latitudes, the energy is carried poleward by transient eddies and by planetary waves which
is represented by the synoptic component of the AHT in
the CLIMBER-2 model (Fig. 3c). As a consequence of
the flattened temperature gradient of the SC3500
experiment, both components of the AHT decrease by
30 to 50% in the SH compared to the CTRL run. In the
NH, the total AHT exhibits smaller changes compared
to the CTRL run which is consistent with the similar
latitudinal temperature gradients found in the NH for
both experiments (Fig. 3a).
The oceanic circulation resulting from the SC3500
experiment is highly asymmetric, with strong sinking
near 60°S reaching a maximum of 35 Sv (Fig. 4a),
considerably stronger than the modern one (21 Sv for
the Atlantic conveyor belt (Petoukhov et al. 2000)). We
find that this thermohaline circulation is rather stable.
Even if the oceanic circulation is perturbed during many
hundred of years to induce Northern Hemispheric
sinking by adding a freshwater flux in the southern high
latitudes (0.1 Sv at 60°S) and a salinity flux in the
northern high latitudes (0.1 Sv at 60°N), the solution
always switches back to sinking in southern high latitudes when the fluxes are removed. In contrast to the
modern system in which maximum OHT is estimated to
be 0.2 to 1 PW greater in the Northern Hemisphere
(Peixoto and Oort, 1992), the boundary conditions used
for the SC3500 run results in much higher heat transport
in the southern hemispheric oceans than in the northern
one (Fig. 4c). The high southward heat flux exhibits a
broad peak of 1.7 PW which spreads from the equator
to nearly 60°S. In the NH, oceanic heat transport shows
a narrower peak of 1 PW in the tropics and reaches
0.5 PW from 30°N to 60°N. This asymmetric pattern
promotes the weaker equator-to-pole temperature gradient simulated in the SH (Fig. 3a) and also the smaller
area of sea-ice in comparison with the NH (Table 1).
As previously explained, one of the critical parameter
in our experiments is the zonal transport of water in
between the three basins at polar latitudes. The circumpolar currents acts as barrier for the overturning
cells and thus may favour the southern sinking mode
due to its larger extent in the north (until 50°) than in the
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Donnadieu et al.: The impact of atmospheric and oceanic heat transports on the sea-ice-albedo instability
stream function induce a modification of the OHT pattern (Fig. 4c) which is still characterised by an asymmetric distribution but with a weaker OHT in the
southern low-to-mid latitudes and with a larger OHT in
the low latitudes than in the SC3500 run. In conclusion,
the asymmetry in the oceanic circulation is controlled
by the continental reconstruction SC but is largely increased by the onset of circumpolar currents. Hence, the
critical parameter, the zonal velocity, has an influence on
the result but is not the only item responsible for the
asymmetric response of the model to the Neoproterozoic
boundary conditions used here.
3.3 Sensitivity to an atmospheric CO2 decrease
Fig. 4 Global vertical overturning stream function simulated in
a the SC3500 run and b the SC3500 NZT run. Negative contours
are dashed. Positive overturning represents clockwise transport.
c Comparison between oceanic annual meridional heat transport
calculated for the CTRL run (black line), and the SC3500 run (grey
line) and the SC3500 NZT run (dashed grey line) (PW)
south (until 60°). The SC3500 run is repeated, with the
only difference being a zonal transport reset at zero
(experiment called ‘NZT’ meaning ‘no zonal transport’.
Southern deep water of the SC3500_NZT run is formed
at a weaker rate of 25 Sv than in the SC3500 run
(Fig. 4a, b) and also penetrates less far away in the NH
but the global asymmetry remains a robust feature of the
model when prescribing no zonal transport between the
three basins. The changes in the meridional overturning
We first analyse the series of simulations with the same
paleogeographic configuration and decreasing level of
CO2 (i.e. SC3500, SC500 and SC200 simulations). In a
second step, the SC90 and SC89 simulations are described with great care because they correspond to the
CO2 threshold levels still yielding an open water solution
in the first case and leading to a global glaciation in the
second one.
The annually averaged mean surface air temperature
distribution versus latitude is shown in Fig. 5a for four
values of the atmospheric CO2 level (3500 ppm,
500 ppm, 200 ppm and 90 ppm). The results indicate
that, as the CO2 concentration decreases, the temperature drops non-uniformly at different latitudes. The decrease of temperature between the SC3500 run and the
SC500 run ranges from 9 °C at the equator to 14–18 °C
at high latitudes. This change in the temperature gradient is, in part, a consequence of the sea-ice-albedo
feedback (Oglesby and Saltzman 1992). The spatial increase of highly reflective surfaces produces a further
decrease in absorbed solar radiation, resulting in a further cooling. This explains why the thermal response was
amplified at high latitudes through this positive feedback
loop. Another feature of the cooling pattern shown in
Fig. 5a is that, although the sea-ice definitively covers
the latitude greater than 60°N and 60°S in the SC500
and SC200 experiments, the high latitudes still undergo a
temperature decrease greater than the low latitudes. The
cooling of the atmospheric column over the high latitudes results in reducing the quantity of atmospheric
water vapour (not shown), which in turn reduce the
greenhouse effect. Figure 5a also illustrates the differential behaviour in between both hemispheres. Indeed,
the cooling is more pronounced in the Southern Hemisphere (particularly at high latitudes) than in the
Northern Hemisphere, which induces a steeper temperature gradient. For example, the cooling between the
SC500 and the SC200 cases reaches more than 14 °C in
the 40–50°S band compared with 10.5 °C in the same
Northern Hemisphere latitudinal band. The sea-ice
cover also exhibits a differential behaviour, it increases
more quickly in the SH than in the NH between the
SC500 and the SC200 cases (see Table 1 and Fig. 5b).
Donnadieu et al.: The impact of atmospheric and oceanic heat transports on the sea-ice-albedo instability
Fig. 5 a Mean annual zonally averaged surface air temperatures
(°C) for the SC3500 run (black line), the SC500 run (black dashed
line), the SC200 run (grey line) and the SC90 run (grey dashed line).
b Same as a but for the mean annual zonally averaged sea-ice cover
fraction (%). c Same as a but for the mean annual meridional AHT
(PW). d Same as a but for the mean annual meridional OHT (PW)
These results may be surprising as we have seen in the
previous section that the NH was colder and with a
larger sea-ice cover than the SH which was characterised
by regions of deepwater formation. In order to fully
understand these differences, we analyse the response of
the atmospheric and the oceanic meridional heat transports which constitute the major diagnostics to understand how these fluids react to the radiative reduction
due to CO2 decrease.
The decrease of CO2 concentration and the subsequent readjustment of the temperature gradient and of
the sea-ice line induce drastic changes in the efficiency of
the atmospheric heat transport (Fig. 5c). In the SH, the
299
peak of the AHT moves from 50°S (SC500) to 30°S
(SC200) and increases in amplitude by 100% (from
SC3500 to SC200). In the NH, the situation is almost the
same but the increase reaches smaller values of about
50%. This is in agreement with the evolution of the
mean zonal temperature mentioned already which
exhibits a steeper temperature gradient in the SH than in
the NH for cold simulations compared with the SC3500
case. Otherwise, the shift of the AHT peak in the SH,
between the SC3500 and the SC200 runs, is strongly
correlated to the migration of the sea-ice line in latitude.
In fact, as the sea-ice edge moves toward low latitudes, it
creates a very steep temperature gradient which drives
the atmospheric heat transport to increase (see the seaice cover, the AHT and the OHT for the SC500 to
SC200 runs in the SH, Fig. 5a–c). In greater details of, it
is interesting to notice the respective role of the synoptic
component at mid-latitudes and of the advective component at low latitudes when the Earth cools (Fig. 6).
Between the SC3500 run and the SC200 run, the sea-ice
front reaches the middle latitudes and then the subtropics. At the same time, the maximum value of the
synoptic component increases and follows the transfer of
the steepest gradient of temperature due to the sea-ice
advance in latitude (Fig. 6a). Conversely, the large increase of the AHT in the tropics is mainly due to the
enhanced advective transport which results from the
Hadley circulation (Fig. 6b).This points toward a critical role of the Hadley circulation acting as a strong
negative feedback once the sea-ice edge approaches the
low latitudes (Lindzen and Farrell 1977; Bendtsen 2002),
thus having an important influence on the stability of the
climate system during a hypothetic decrease of the
atmospheric CO2 levels.
As the CO2 level decreases from 3500 ppm to
200 ppm (the last simulation will be analysed later), the
oceanic heat transport exhibits a large decrease in the
SH whereas in the NH the ocean heat transport increases in the tropics and decreases in the mid-to-high
latitudes (Fig. 5d). In both hemispheres, at mid-to-high
latitudes, the OHT decrease is in part a result of the
equatorward migration of the sea-ice but to understand
the large reduction of the OHT around 30°S (see e.g. the
evolution between the SC3500 and the SC500 simulations) and the progressive increase of the OHT around
20°N requires us to analyse in more detail the evolution
of the oceanic circulation in the model. Figures 7a, b
displays the mean annual meridional overturning stream
function for the SC500 and SC200 simulations.
Within the SH, firstly, the sea-ice migration clearly
induces equatorward displacement of the location of the
oceanic deep convection and thus results in a decrease of
the OHT poleward of the deepwater formation areas
(Figs. 5b, d and 7a, b). Secondly, as the CO2 level decreases, the intensity of the ocean circulation is reduced
in the SH from 35 Sv in SC3500 to 20 Sv in SC200
associated with the OHT diminution seen north of the
area of deepwater formation (Figs. 5 and 6). Several
factors act as a potential inhibitor of the overturning.
300
Donnadieu et al.: The impact of atmospheric and oceanic heat transports on the sea-ice-albedo instability
Fig. 6 a Synoptic component of the AHT for the SC3500 run
(black line), the SC500 run (black dashed line), the SC200 run (grey
line) and the SC90 run (grey dashed line). The vertical lines show
the lowest latitudes of the mean annual sea-ice cover for the SC500
run (black dashed line), the SC200 run (grey line) and the SC90 run
(grey dashed line). Note the synchronous migration of the sea-ice
and of the maximum synoptic heat transport. b Same as a but for
the advective component
Firstly, the decrease of the tropical zonal averaged
temperature and the equatorward migration of the seaice (note that the temperature at the ice margin remains
close to the freezing point –1.8 °C) reduce the equator to
pole ocean temperature gradient. The thermal forcing on
the overturning cell, which is primarily due to the temperature difference between the ‘warm-water’ and the
‘cold-water’ sphere, is then reduced and can explained,
in part, the decrease of the overturning. Secondly, the
melting of the sea-ice delivers a larger freshwater flux in
the front of the area of deepwater formation when the
CO2 decreases (Fig. 8b), this can potentially reduce the
overturning. This feature occurs in response to the enhanced AHT brought into the front of the sea-ice margin and also to the enhanced westerly winds that force
the sea-ice to be deriven equatorward (from the SC3500
to the SC200 runs, the wind stress increases by 26%
between 60°S and 30°S and by 23% between 60°N and
30°N). The quantity of melted sea-ice is affected by both
mechanisms. This transport of sea-ice from mid-to-low
latitudes due to the westerly winds results in an inflow of
latent heat of melting which may infer the weak poleward OHT of less than 0.2–0.3 PW seen below the
Fig. 7 Global vertical overturning stream function simulated in
a SC500 run, b the SC200 run and c the SC200_SIS (see the text for
the abbreviations). Negative contours are dashed. Positive overturning represents clockwise transport. Note the existence of a
secondary cell located below the ice between 40°S and 60°S. This
overturning cell is mainly due to the driving effect of the water
sinking in the front of the sea-ice margin. Indeed, the activity of this
cell increases when the main anticlockwise cell located north of
40°S increases (compare the SC200 and the SC200_SIS runs)
sea-ice in Fig. 5d. In order to ascertain the significance
of the freshwater-feedback on the overturning, the suite
of SC experiments has been re-run by removing the effect from melting or formation of sea-ice on the surface
salinity (experiments known as SIS meaning sea ice
sensitivity). In Fig. 7c, we have plotted the overturning
function for the SC200_SIS run. In comparison with the
SC200 run, the deep water formation is clearly more
active when the freshwater-feedback is removed. However, although the sensitivity of the model to the radiative forcing decreases as the CO2 threshold to achieve a
hard snowball Earth is 74 ppm in the SC_SIS runs
(89 ppm in the SC runs), this effect remains limited.
Donnadieu et al.: The impact of atmospheric and oceanic heat transports on the sea-ice-albedo instability
Fig. 8a–c Freshwater fluxes in the ocean realm for the SC3500 run
(black line), the SC200 run (black dashed line) and the SC90 run
(grey line). a The mean annual global balance (mm/day). b Fluxes
due to the melting and/or to the formation of the sea-ice (mm/day).
c Fluxes due to the atmospheric component, i.e. the sum of the
precipitation and the runoff minus the evaporation averaged over
the ocean realm (mm/day)
Conversely, experiments with decreased CO2 levels
have strengthened the northern meridional upper circulation cell (Figs. 7a, b) which induces enhanced ocean
heat transport in the northern tropics (Fig. 5d). In
addition and more importantly, along the cooling sector,
the large southward cross-equatorial transport of heat
which characterised the SC3500 run is reduced while at
the same time the OHT at 20°N increases almost linearly
(Fig. 5d). This demonstrates that the reduction of the
cross-equatorial transport of deep water from the
southern overturning cell, seen throughout the cooling
(Figs. 4a and 7a, b), induces the readjustment of the
OHT between both hemispheres until the OHT reached
a nearly symmetrical solution around the CO2 level
200 ppm (Fig. 5d).
We have also investigated the sensitivity of these results to the prescribed zonal velocities. Surprisingly, the
OHT for the SC500_NZT and for the SC200_NZT runs
are almost the same than those simulated in the runs
with the zonal transport (this is why OHT from
SC500_NZT and SC200_NZT runs are not shown as
they are merged with the OHT plotted in Fig. 5d). In
fact, once the area of deep water formation moved to
mid-latitudes (Fig. 7), the overturning cells undergo no
more the influence of the circumpolar currents.
In comparison with the adjustment of the atmosphere
and ocean to a CO2 decrease, it clearly appears that the
301
atmosphere acts as a strong negative feedback while the
response of the ocean is somehow more complex (negative feedback in the Northern Hemisphere and positive
feedback in the Southern Hemisphere) depending on the
thermohaline circulation, but also on the coupled response of the atmosphere-ocean system via the enhanced
dynamic circulation of the atmosphere and the steeper
latitudinal temperature gradient generated in cold
climates.
The simulation at a CO2 level of 90 ppm that still
yields an ice-free tropical ocean is now investigated. At
high to mid-latitudes, the AHT decreases while the
strengthening of the Hadley cell (18% in term of mass
transport in the SH between the SC200 run and the
SC90 run) promotes an increase of the heat transfer in
the low latitudes (Fig. 5c). The OHT decreases at all
latitudes and in both hemispheres (Fig. 5d). The main
reason is related to the combined effects of (1) the
cooling of the equatorial surface waters [mean tropical
SST is around 9 °C (4 °C) in the SC200 (SC90) experiment] and (2) the advance of the sea-ice line which act
together to reduce the zonal gradient of temperature
and, thus, the transport of equatorial water via the
overturning cell as already explained. In addition but
less importantly, once the sea-ice has reached the poleward limit of the tropics, it discharges large freshwater
fluxes (Fig. 8b) in the low latitudes. The salinity flux
arising from the atmospheric component located over
the tropics also exhibits a large reduction between the
SC200 run and the SC90 run (Fig. 8c). Both mechanisms
contribute to the decrease of the global salinity flux over
low latitudes (Fig. 8a) and result in a diminution of
convective activities (not shown) and of the OHT.
At a CO2 level of 89 ppm (global glaciation), the
transition from a soft to a hard snowball Earth takes a
few hundreds of years (Fig. 9a). Initially the mean global temperature decreases slightly (over 250 years), the
surface tropical ocean loses its heat via the large sensible
fluxes occurring in the front of the ice margin and cools
by 2.5 °C [from 4 °C (SC90) to 1.5 °C] but the sea-ice
margin remains at around 30° of latitudes in both
hemisphere (Fig. 9b). The OHT goes on to decrease;
although the atmospheric heat transfer toward the
tropical latitudes remains important, it also begins to
decrease at subtropical latitudes (Fig. 9c, d). In a second
step, the transition to a globally ice-covered ocean takes
a few years (23 years in this simulation). At this time, the
response of the Hadley circulation is interesting in terms
of feedback. As long as the sea-ice does not reach the
tropics, the Hadley cells work in the opposite direction
to the sea-ice advance by homogenising the temperatures
in the tropics. However, once the sea-ice achieves a
tropical location, the cold surface air temperatures at
these latitudes will be transported to the equator by the
lower meridional branch of the Hadley cell (Bendtsen
2002). As a consequence, the temperatures in the tropics
approach freezing point and lead the atmospheric negative feedback to shut down. This can be seen in
Fig. 9a–c where the modification of the AHT at the
302
Donnadieu et al.: The impact of atmospheric and oceanic heat transports on the sea-ice-albedo instability
model years 250 and 268 is combined with the migration
of the ice margin, so that synchronously the OHT collapses. After this, the duration required by the sea-ice to
reach the equator is only related to the ocean thermal
buffering capacity (Bendtsen and Bjerrum 2002). Once
the sea-ice entirely covered the oceans (the hard snowball Earth state), the OHT becomes ineffective and the
AHT is rather weak with maximum values of 1–1.5 PW
(Fig. 9). The mean global temperature reaches –42 °C.
3.4 Sensitivity to the paleogeography
The paleogeography used in the DC experiments represents the continental configuration after the dislocation of the Rodinia super continent. The tropical land
surface area remains approximately unchanged, about
70% in both cases as well as the mean zonal land surface
area differs little in each configuration. Thus, it is not
surprising that the model’s sensitivity to reductions in
atmospheric CO2 is slightly the same between 5000 ppm
and 500 ppm for both paleogeographies (Fig. 2). However a divergence appears below 280 ppm where the
cooling trend of the mean temperature becomes steeper
in the DC case than in the SC case. Thus, the DC
experiments plunge into a snowball Earth when the CO2
threshold value of 149 ppm is reached whereas the SC
experiments require an additional decrease of 60 ppm to
reach a completely ice covered state. In order to explore
the mechanisms inducing this divergence, the SC and
DC experiments at 280 ppm, 200 ppm and 150 ppm are
investigated. Figure 10 illustrates the mean annual
AHT, OHT and SST (sea surface temperature) for latitudes between 45°S and 45°N. For the CO2 level
280 ppm (not shown) and for the three terms, the curves
from the SC and DC simulations are almost merged. For
the CO2 level 200 ppm, in the SH, the oceanic heat
transport reaches smaller values in the DC case than in
the SC case (Fig. 10b) while in the NH, they are similar
in amplitude. In the Southern Hemisphere, the equatorward extent of sea-ice is larger in the DC case than in
the SC case (Fig. 10c). For the CO2 level 150 ppm, the
decrease of the oceanic heat transport in the DC case
compared to the SC case is clear at all latitudes
(Fig. 10d). The sea-ice line reaches the poleward rim of
the 4 tropics and the SSTs are 2–3 °C colder in the DC
case (Fig. 10e). At the same time, the atmosphere acts,
as already shown as a negative feedback by increasing
the heat transport at low latitudes, and exhibits small
differences in between the two continental configurations (Fig. 10a, d). The last step to induce climatic
instability in the DC case is similar to the mechanism
already explained for the SC case: once the sea-ice advances in to the tropics and induces air surface temperatures below 0 °C, the tropical temperature being
homogenised by the Hadley circulation also decreases
and then leads to a large reduction of the AHT.
The possible effects of the oceanic circulation on the
enhanced sensitivity of the DC simulations to radiative
Fig. 9 a Global annual mean surface-air temperatures (°C) versus
year of model run for the SC89 experiment. b Geographical
distribution of the sea-ice 160 years, 250 years and 268 years after
the beginning of the SC89 simulation. c Atmospheric annual
meridional heat transport for (1) the years of the SC89 run, 160
(black line), 250 (dashed black line) and 268 (grey line) and (2) once
equilibrated (dashed grey line) (PW). d Oceanic annual meridional
heat transport for (1) the years of the SC89 run, 160 (black line),
250 (dashed black line) and 268 (grey line) and (2) once equilibrated
(dashed grey line) (PW)
forcing are now explored. Firstly, up to a level of
280 ppm, the DC simulations are also characterised by
an asymmetric oceanic circulation associated deep water
sinking with strong mainly located in the SH but the
meridional overturning stream function is substantially
more sluggish than in the SC simulations (Fig. 11a),
which leads to colder deep ocean in the DC case (about
0.5 °C less at 500 ppm). Thus, the ocean thermal buffering heat capacity as the sea-ice approached the low
latitudes would be less important in the DC case.
Donnadieu et al.: The impact of atmospheric and oceanic heat transports on the sea-ice-albedo instability
303
Fig. 10 a Comparison between
atmospheric annual meridional
heat transport calculated for the
SC200 run (black line) and the
DC200 run (grey line) (PW).
b Same as a but for the oceanic
heat transport (PW). c Same as
a but for the mean zonal sea
surface temperature (°C).
d Comparison between
atmospheric annual meridional
heat transport calculated for the
SC150 run (black line) and the
DC150 run (grey line) (PW).
e Same as d but for the oceanic
heat transport (PW). f Same as
d but for the mean zonal sea
surface temperature (°C)
Second and more importantly, below 280 ppm, the
oceanic circulation differs greatly between the SC runs
and the DC runs (Fig. 11b–c). The SC simulations are
characterised by a strong sinking, located in front of
the sea-ice margin, whereas the circulation of the DC
simulations switches into a new equilibrium where the
main location of the oceanic deep convection is displaced below the sea-ice around 60°S. Thus, on one
hand, the high to mid latitude intermediate and deep
cold waters are isolated from the ‘hot’ low-latitude
water; on the other hand, the DC large-scale advective
circulation mixes the high-latitude waters with the lowlatitude waters in the SH. These changes are in part
due to the modification of the surface freshwater flux
forcing (Fig. 12). Between the DC280 run and the
DC150 run (Fig. 12b), freshwater flux reaches negative
values at 70–50°S (equivalent to a salinity flux) while it
increases at 40–30°S, hence these changes promote an
increase (decrease) of the convective activity at 70–
50°S (40–30°S). This is in agreement with the change
in the mode of oceanic circulation seen already above.
The surface freshwater flux differences between the
SC280 run and the SC150 run are less important and
thus, do not promote a shift of the thermohaline
mode.
The comparison of the behaviour of the ocean circulation in the SC and DC configurations points out
that in the DC simulations, because of changes in the
hydrologic cycle, there is a possibility to shift to another
equilibrium that acts as a powerful feedback to produce
the global glaciation which is achieved for a higher CO2
level than in the SC simulations. Indeed, the modified
thermohaline circulation in the DC150 run generates
upwelled cold waters in the tropics in agreement with the
decrease of the SST (Fig. 10f). In addition, as a consequence of the decreased meridional overturning circulation around 35°S (Fig. 11b), southern OHT is reduced
and allows an equatorward migration of the sea-ice edge
that is more important in the DC case than in the SC
case (Fig. 10d–e).
4 Discussion
There have been several studies on the initiation of the
snowball state using atmospheric GCMs coupled to a
mixed-layer ocean (Jenkins and Smith 1999; Chandler
and Sohl 2000; Hyde et al. 2000; Donnadieu et al. 2003)
or using the AOGCM FOAM (Poulsen et al. 2001,
2002). A variety of paleogeographies has been used in
these studies which does not allow a direct comparison
of our results with all of these experiments. Nevertheless,
the studies of Jenkins and Smith (1999) and of Poulsen
et al. (2002) used a single idealised supercontinent centered on the equator that is broadly comparable to our
SC reconstruction in terms of latitudinal distribution.
The pCO2 thresholds, 340 ppm or larger, obtained by
the uncoupled atmospheric GCM studies (Jenkins and
Smith 1999; Donnadieu et al. 2003) is greater than those
obtained here, 90 ppm and 150 ppm, which demonstrates the important role played by ocean dynamics in
delaying the onset of the sea-ice instability in the
CLIMBER-2 model. In comparison with the FOAM
experiments, CLIMBER-2 exhibits a greater climatic
sensitivity to reductions in CO2 contents, somehow as a
result of the existence of a thermodynamic sea-ice model
that is not included in the FOAM model. Indeed, the
effect of the sea-ice on the freshwater surface fluxes (and
thus on the oceanic circulation) becomes important
during the cooling in our experiments. More fundamentally, the value of the critical concentration of CO2
which leads to the collapse may be significantly affected
by the GCM’s sea-ice albedo values but also (and it is a
concern never addressed in the Neoproterozoic studies)
by the land surface albedo values which shows large
difference over the desert areas in between the GCMs
(see (Bonfils et al. 2001)). Moreover, the short duration
(60 years) of the experiments presented in Poulsen et al.
(2002) does not allow them to account for changes in
ocean dynamics but only to the fast response of the
surface layer in the tropical area. The fact that the CO2
304
Donnadieu et al.: The impact of atmospheric and oceanic heat transports on the sea-ice-albedo instability
Fig. 12 a Mean annual surface freshwater flux (positive values for
an input of freshwater into the ocean) for the SC280 run (black line)
and the SC150 run (dashed black line). b Same as a but for the
DC280 run (black line) and the DC150 run (dashed black line)
Fig. 11 Comparison between global vertical overturning stream
function in a the DC500 run, b the DC150 run and c the SC150 run.
Negative contours are dashed. Positive overturning represents
clockwise transport
threshold found for the SC and DC simulations are in
the range of previous simulations demonstrates the
ability of an EMIC to cope with deep-time paleoclimate
studies. We show that accounting for ocean and atmosphere dynamics certainly reduces this threshold.
Moreover, we also show that it depends on the paleogeography.
Our attempt was also to improve our understanding
of the mechanisms that participate in or that counteract
the initiation of the climate instability. In terms of
feedbacks, our study clearly demonstrates that the
meridional atmospheric heat transport plays the most
important role by increasing the energy brought in to the
front of the sea-ice margin throughout the cooling. We
have also shown the critical role of the atmospheric
circulation via the Hadley cells to counteract the
migration of the sea-ice inside the tropics and thus their
influence on the onset of the sea-ice-albedo instability.
This negative feedback shows few changes between a
supercontinent setting or a fragmented continental
configuration that, however, is characterised by a similar
latitudinal distribution of the land area. Future studies
are planned to assess the response of the low-latitude
atmospheric circulation in a geographic configuration in
which most of the tropics would be covered by oceans
such as the configuration thought to typify the Varangian-Vendian episodes (620–580 Ma) (Cawood and
Nemchin 2001; Meert 2003).
The role played by the ocean is more complex; in
term of feedbacks, no clear picture emerges from the
CO2 sensitivity experiments. The fact that the simulation
with a high CO2 level depicts a large asymmetry between
the Northern and Southern Hemispheres, with deep
convection occurring strongly and essentially in SH high
latitudes is very important to explain the opposite
behaviour of the ocean heat transport in both hemispheres when the CO2 values decrease. In short,
throughout the cooling of the Earth, the OHT is continuously readjusted between both hemispheres until it
reached a nearly symmetrical situation around the CO2
level 200 ppm (for the SC simulations). In the SH, the
decrease of the OHT is mainly due to the changes of the
intensity of the thermohaline circulation with a weaker
vertical mixing contributing to a greater extent of the
sea-ice margin. Using a three-dimensional ocean model
(the GFDL modular ocean model) forced by output
from AGCM experiments, Bice et al. (2000) found the
same behaviour as they demonstrated that, from 55 Ma
to 14 Ma, the Earth has experienced an overall decrease
in SH poleward OHT and an increase in NH transport
due to subtle paleogeographic changes. Below 200 ppm,
Donnadieu et al.: The impact of atmospheric and oceanic heat transports on the sea-ice-albedo instability
the general trend is modified and the OHT undergoes an
overall decrease in both hemispheres because of the decrease of the thermal forcing on the overturning cells due
to the combined effects of the tropical SST cooling and
of the sea-ice advance which reduce the temperature
zonal gradient. This behaviour of the OHT emphasises
the inability of the ocean to act as a negative feedback
once the sea-ice reaches latitudes 30°. However, the
oceanic circulation can influence the CO2 threshold below which the climate instability takes place. The results
of the DC simulations demonstrate that, via changes in
the continental configuration, the oceanic circulation is
modified and enhances the sensitivity of the Earth to a
CO2 reduction (from 90 ppm in the SC experiments to
150 ppm in the DC experiments) but it does not seem
that dynamic oceanic processes can prevent the onset of
the ice-albedo instability.
To summarise, a new set of a coupled atmosphereocean model simulations with variable paleogeographies
provides additional support for a possible snowball
Earth scenario during the Neoproterozoic period. Our
results also suggest that a CO2 level of 1500 ppm is required to maintain the mean global temperature above
10 °C with a 6% decrease of the solar constant. This
conditions could have existed before the Rodinia
breakup as there are no clues suggesting glacial periods
between 2.5 Ga and 800 Ma. Thus, our contribution
poses questions for geochemical modellers as one needs
to find a mechanism which consumes more than
1000 ppm of atmospheric CO2 in order to drive the
Earth in to favourable climatic settings for a snowball
glaciation.
Acknowledgements Y.D. wishes to thanks D. Paillard for many
stimulating discussions in the field and also J. Meert who kindly
provided the maps. This work is supported by the French program
ECLIPSE. Computing was carried out at LSCE, supported by the
CEA.
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