The Eocene–Oligocene transition (EOT), which took place approximately 34 Ma ago, is an
interval of great interest in Earth's climate history, due to the inception
of the Antarctic ice sheet and major global cooling. Climate
simulations of the transition are needed to help interpret proxy data,
test mechanistic hypotheses for the transition and determine the climate
sensitivity at the time. However, model studies of the EOT thus far typically
employ control states designed for a different time period, or ocean
resolution on the order of 3∘. Here we developed a new higher
resolution palaeoclimate model configuration based on the GFDL CM2.1 climate
model adapted to a late Eocene (38 Ma) palaeogeography reconstruction. The
ocean and atmosphere horizontal resolutions are 1∘× 1.5∘
and 3∘× 3.75∘ respectively. This represents a
significant step forward in resolving the ocean geography, gateways and
circulation in a coupled climate model of this period. We run the model under
three different levels of atmospheric CO2: 400, 800 and 1600 ppm. The model
exhibits relatively high sensitivity to CO2 compared with other recent
model studies, and thus can capture the expected Eocene high latitude warmth
within observed estimates of atmospheric CO2. However, the model does
not capture the low meridional temperature gradient seen in proxies.
Equatorial sea surface temperatures are too high in the model
(30–37 ∘C) compared with observations (max 32 ∘C), although
observations are lacking in the warmest regions of the western Pacific. The
model exhibits bipolar sinking in the North Pacific and Southern Ocean, which
persists under all levels of CO2. North Atlantic surface salinities are
too fresh to permit sinking (25–30 psu), due to surface transport from the
very fresh Arctic (∼ 20 psu), where surface salinities
approximately agree with Eocene proxy estimates. North Atlantic salinity
increases by 1–2 psu when CO2 is halved, and similarly freshens when
CO2 is doubled, due to changes in the hydrological cycle.
Introduction
The Eocene–Oligocene transition (EOT), which occurred approximately 34 Ma ago, was one of
the major climate transitions in the Cenozoic era. The EOT marks the shift
from the so-called “greenhouse world” to the “icehouse world”, during which
time the first semi-permanent ice sheets formed on Antarctica, and major
changes occurred in flora and fauna across the globe reflecting a shift to
colder or drier conditions (Miller et al., 1991; Coxall and Pearson, 2007;
Dupont-Nivet et al., 2007). Yet the causes of the
transition and its global impacts are still not fully understood,
particularly regarding the role of ocean circulation. The prevailing
proposed mechanisms for the transition are: a decline in atmospheric
CO2 below a critical threshold (DeConto and Pollard, 2003),
and the opening of Southern Ocean gateways causing the thermal isolation of
Antarctica (Kennett, 1977); with orbital forcing playing a key role as trigger and pacemaker of
ice growth on Antarctica (Coxall et al., 2005). In all
cases, ice–albedo feedbacks play a major role in enabling the beginning of
land-ice formation to grow into a continental ice sheet. Palaeogeography has
also been found to partly control the sensitivity of these mechanisms, as
illustrated by Lear and Lunt (2016).
A long-term decline in atmospheric CO2 has been found to be a major
driver of Eocene cooling (Anagnostou et al., 2016), and a
plausible trigger of the EOT glaciation (DeConto and Pollard,
2003). Hereafter we refer to “atmospheric CO2” as simply CO2.
This mechanism is appealing in its simplicity, however the threshold
required to trigger the glaciation is uncertain. Gasson et
al. (2014) found the glaciation threshold to be within a range of 560 to 920 ppm using an intercomparison of late Eocene climate models to force an ice
sheet model. This large range was primarily due to the spread in climate
sensitivity and seasonality of the models, of between 2.5 and
4.1 ∘C per doubling of CO2 under modern conditions.
Observations of CO2 from just prior to the EOT are similarly uncertain,
with estimates from carbonate microfossils providing an uncertainty range of
450 to 1500 ppm with a central estimate of 760 ppm (Pearson
et al., 2009). Strong CO2 forcing has had considerable success in
explaining the very warm climate of the early Eocene (Huber and
Caballero, 2011). The subsequent increase in seasonality and the meridional
temperature gradient from the early to late Eocene supports the idea that
CO2 forcing had a primary role in Eocene to Oligocene climate change
(Eldrett et al., 2009).
The evolution of the oceanic meridional overturning circulation (MOC) around
the EOT remains a topic of debate. Global compilations of benthic
foraminiferal stable isotopes of δ13C and δ18O
have been interpreted as indicating that the deep ocean was largely
homogeneous during the early to middle Eocene (Cramer et al.,
2009). This has been used to infer a predominance of Southern Hemisphere
high latitude deep-water formation. Subsequent divergence of inter-basin
benthic isotopic gradients creating oceanic heterogeneity has been linked to
onset of North Atlantic intermediate water formation in the late Eocene, and
the beginning of a more mature North Atlantic deep water formation in the
early Oligocene (Borrelli et al., 2014). In those studies,
the proposed forcing mechanism of North Atlantic deep water is the opening
of Southern Ocean gateways. This notion rests on the modern climate
interpretation that, in addition to diapycnal mixing in the interior ocean
(Munk and Wunsch, 1998), wind-driven upwelling in the Antarctic Circumpolar Current (ACC) is a
major driver of the deepwater cell presently associated with the Atlantic
MOC (Toggweiler and Samuels, 1995). However, transport by
mesoscale eddies counter the wind-driven upwelling in the ACC and the
strength of the MOC (Marshall and Speer, 2012). Recent modelling
suggests that deepening of the Greenland–Scotland Ridge may be a trigger for
the onset of North Atlantic sinking (Abelson and Erez, 2017), and
the resulting changes in heat transport could also have important effects on
the period's climate.
In contrast to the North Atlantic sinking hypothesis, Nd isotope
distributions suggest that North Pacific sinking occurred all the way from
the Cretaceous through to the Miocene (Ferreira et al., 2018).
Nd tracer-enabled ocean simulations that are able to reproduce the observed
pattern of Eocene Nd have vigorous bipolar Pacific sinking, with southern
rather than North Pacific bottom water dominating (Thomas et
al., 2014). Furthermore, idealised climate model simulations show that salt advection
feedbacks create competition between North Pacific and North Atlantic deep
water formation (Wolfe and Cessi, 2014), so that the
onset of North Atlantic sinking may imply a shutdown of North Pacific
sinking and vice versa. It remains unclear if such a switch occurred at the
EOT, and its potential triggers have not been fully explored.
Modelling studies investigating the effect of the ACC on global climate have found that opening the Drake Passage
leads to cooling of the Southern Ocean and a reduction in poleward heat
transport towards Antarctica (Cox, 1989; Sijp et al., 2009;
Sijp and England, 2004). These studies used either ocean-only or
intermediate complexity climate models lacking dynamic atmospheric
feedbacks. The Antarctic response of fully coupled climate models to Drake
Passage opening is less clear, since the opening of a circumpolar gateway
does not guarantee a strong ACC under late Eocene
palaeogeography (Zhang et al., 2010), or under the high CO2
(Lefebvre et al., 2012) prevalent in the late Eocene.
Furthermore, there is uncertainty of up to tens of millions of years over
when a deep gateway first existed in Drake Passage
(Livermore et al., 2005; Barker et al., 2007),
so that the inception of the ACC cannot be easily pinpointed to the EOT. The
opening of a deep Tasman Seaway has also been suggested to have occured around the
EOT and caused Antarctic glaciation (Stickley et al., 2004). A
model study of this opening found that the gateway change had a limited
effect on Antarctic surface climate (Sijp et al., 2011).
However, the gateway change may explain ∼ 3 ∘C of
deep ocean cooling (Sijp et al., 2014), in agreement with
late Eocene cooling found in deep sea foraminifera
(Bohaty et al., 2012). Part of the CO2 decline at
the EOT may have been caused by the opening of Southern Ocean gateways
(Elsworth et al., 2017), due to feedbacks in silicate
weathering. This presents a possible way to reconcile the CO2 forcing
and gateway hypotheses.
Bathymetry and topography of the late Eocene, adapted
from Baatsen et al. (2016). Both figures are plotted using a cell fill
method, illustrating the resolution of the grid cells in each case. Due to
the difference in resolution between the ocean and the atmosphere, coastal grid
cells in the atmosphere typically contain a fraction of both land and ocean.
A common shortcoming of previous EOT simulations is the unrealistic
representation of topography (by “topography” we mean both land topography
and ocean bathymetry). Many studies investigating gateway effects have used
modern topography for their control simulation
(Sijp et al., 2009; Yang et al., 2014; Fyke et al., 2015; Elsworth et al., 2017; England et al., 2017) and created EOT-like perturbations by closing Drake
Passage and/or opening the Panama Seaway. While this method of changing
basin geometry is appealing in that it illuminates specific gateway effects,
it implicitly ignores several features of the late Eocene or early Oligocene
palaeogeography that are potentially crucial to the ocean circulation. These
include particularly the closed Pacific–Arctic Bering Sea seaway, the
narrower Atlantic Basin, the shallow connections between the Atlantic and
the Arctic, and the connections of the Tethys Ocean with the Indian and
Arctic oceans (Fig. 1). Other modelling studies
of the EOT have been configured with early Eocene (∼ 55 Ma)
palaeogeography. These have particularly focused on the effects of changing
CO2 forcing from the Eocene to the Oligocene
(Eldrett et al., 2009; Liu et al., 2009),
the impact of an ice sheet on the atmosphere and ocean circulation
(Goldner et al., 2013, 2014) and the impact of opening
Southern Ocean gateways (Goldner et al., 2014). While early
Eocene topography is a much closer starting point to late Eocene topography
than modern, early and late Eocene topographies differ significantly due to
∼ 20 Ma of tectonic evolution. This includes the gradual
opening of the Tasman Gateway and Drake Passage, and the widening of the
Atlantic.
Some recent climate model studies have used reconstructions of late Eocene
and early Oligocene topography for their control runs
(Zhang et al., 2010; Lefebvre et al., 2012; Ladant et al., 2014; Kennedy et al., 2015; Lunt et al., 2016). Kennedy et al. (2015)
found that when an ice sheet is placed over Antarctica, the Pacific sector
of the Southern Ocean surface warms significantly when the Tasman Seaway is
constricted (as in the late Eocene), whereas when the Tasman Seaway is wider
(as in the early Oligocene), there is instead a weak cooling. These results
contrast with Goldner et al. (2014), who found a strong
cooling of the Southern Ocean in response to Antarctic glaciation, albeit
with an early Eocene palaeogeography. Thus, uncertainty remains regarding
whether the opening of Southern Ocean gateways caused major cooling, or
whether the glaciation itself was the primary driver of Southern Ocean
temperature change. A summary table of previous EOT climate simulations is
shown in Table 1.
A non-exhaustive summary of previous EOT climate simulations,
including their ocean and atmosphere horizontal resolutions, age of
palaeogeography and sinking regions. EBM represents a 2-D energy balance
model. All are coupled climate models except the MITgcm
(Thomas et al., 2014) and POP2.1 (Baatsen
et al., 2018b) simulations, which are included because they specifically
investigate the meridional overturning circulation (MOC) around the EOT.
n/a means “not applicable”.
ModelOcean resolutionAtmos. resolutionAge ofpalaeogeographySinking regionsReferenceCCSM1.41.8∘× 3.6∘3.75∘Early Eocene(Sewall et al., 2000)Southern Ocean andN. AtlanticHuber andSloan (2001);Huber et al. (2004)CCSM31∘× 2.8∘3.75∘Early Eocene (Sewall et al., 2000)Southern OceanEldrett etal. (2009); Liu etal. (2009)CESM1.01.8∘× 3.6∘3.75∘Early Eocene (Sewall et al., 2000)Southern OceanGoldner etal. (2014)CESM1.0.51∘2∘Late Eocene 38 Ma(Baatsen et al., 2016)S. PacificonlyBaatsen etal. (2018a);in reviewCM2Mc2.5∘× 0.6–3∘3∘× 3.75∘Modern + gateway perturbationsSouthern Ocean andN. AtlanticYang et al. (2014); Elsworth etal. (2017)COSMOS1.8∘× 3∘3.75∘Early Miocene 20–15 Ma Herold et al.,2008) + gatewayperturbationsSouthern Ocean andN. AtlanticStärz et al. (2017)FOAM1.4∘× 2.8∘4.5∘× 7.5∘Early Eocene (Zhanget al., 2010)Not statedZhang et al. (2010)FOAM1.4∘× 2.8∘4.5∘× 7.5∘Middle Oligocene(Scotese, 2001)Southern OceanLefebvre etal. (2012); Ladant et al. (2014)HadCM3L2.5∘× 3.75∘2.5∘× 3.75∘-38 to 34 MaRupelian -34 to 30 MaPriabonian (GETECH – propri-etary)Southern Ocean and N. AtlanticInglis et al. (2015); Kennedy etal. (2015)MITgcm (ocean only)4∘n/aEarly Eocene (Sewall et al., 2000)S. Pacificand N. PacificThomas etal. (2014)POP2.1 (ocean only)1∘n/aLate Eocene 38 Ma(Baatsen etal., 2016)S. Pacific or N. PacificBaatsen etal. (2018b)UVic1.8∘× 3.6∘EBMModern + gateway perturbationsSouthern Ocean andN. AtlanticFyke et al. (2015)UVic1.8∘× 2.4∘EBMEarly Eocene (Huber et al., 2004)Southern OceanSijp et al. (2011)GFDL CM2.11∘× 1.5∘3∘× 3.75∘Late Eocene 38 Ma(Baatsen et al., 2016)Southern Ocean andN. PacificThis study
Model studies of the EOT, including those mentioned above, generally employ
very low – typically 3∘ – horizontal ocean resolution. This is
often dictated by computational constraints, since the deep ocean takes
thousands of years to equilibrate (Danabasoglu et al., 1996). Furthermore,
uncertainties in the reconstructed palaeogeography often place limitations on
the fidelity of the relevant topographic features. We are interested in
improving the ocean resolution for two reasons. First, the Arakawa B-grid
used ubiquitously in palaeoclimate model studies requires narrow ocean
straits to be at least two grid cells wide, in order to have a non-zero
velocity grid point. If a model uses 3∘ resolution, then the minimum
width of a strait is 6∘, corresponding to 670 km in the meridional
direction. In the zonal direction, 6∘ of distance varies from 670 km
at the equator to 330 km at 60∘ latitude. These distances are large
enough that key EOT ocean gateways, such as Drake Passage and Arctic gateways
will hardly be resolved. Second, there is evidence that gateway effects are
enhanced at higher resolution (Viebahn et al., 2016), especially in
controlling the heat transport of boundary currents. For example, moving from
1∘ to 0.25∘ (eddy-permitting) resolution in the modern-day
climate results in an enhanced western boundary current heat transport
(Delworth et al., 2012), and this effect can cause substantially lower Arctic
sea ice coverage (Hutchinson et al., 2015; Kirtman et al., 2012). Moving from
3 to 1∘ ocean resolution, while still not permitting eddies, would
improve the representation of boundary currents and their associated impacts
on gateway transitions.
This study presents a new fully coupled climate simulation of the late Eocene
using order 1∘ ocean resolution and a state-of-the-art
palaeotopographic reconstruction. We apply the late Eocene (38 Ma)
reconstruction of Baatsen et al. (2016) to generate topography for the GFDL
CM2.1 model, as shown in Fig. 1. This ensures that global tectonic evolution
relevant to the late Eocene is included. Within this framework, we explore
the climate sensitivity to CO2 perturbations, namely using 400, 800
and 1600 ppm (approximately 1.4×, 2.9× and 5.7×
pre-industrial CO2 respectively). We also explore the sensitivity
to changing the parameterized ocean vertical mixing from the standard
Bryan–Lewis scheme to a tidal mixing scheme, where the diffusivity is set to
a constant background value and enhanced near the bottom. We examine the
impact of these changes on the oceanic MOC and global climate.
Model descriptionModel components
This study uses a modified version of the coupled climate model GFDL CM2.1
(Delworth et al., 2006), using boundary conditions for the
late Eocene. The ocean component is updated to the modular ocean model (MOM)
version 5.1.0, while the other components of the model are the same as in
CM2.1, namely Atmosphere Model 2, Land Model 2 and the Sea Ice Simulator. In
the ocean and sea ice components, the horizontal grid is modified to have a
resolution of 1∘ latitude × 1.5∘ longitude. We use a
tripolar grid as shown in Fig. 7 of Murray (1996),
with a regular latitude–longitude mesh south of 65∘ N, and a
transition to a bipolar Arctic grid north of 65∘ N. The poles then
lie over North America and Siberia to avoid convergence of meridians in the
Arctic Ocean. This enables the model to simulate Arctic Ocean velocities
without damping in the vicinity of the North Pole. Unlike in CM2.1, there is
no refinement of the latitudinal grid spacing in the tropics. The ocean
retains the original 50 vertical levels with the same grid spacing as in
CM2.1. The atmospheric horizontal grid resolution is 3∘× 3.75∘ , with 24 vertical levels. The atmosphere grid is identical
to that used in CM2Mc (Galbraith et al., 2010). This choice of
grid resolutions gives good load balancing between the atmosphere and ocean
components, while enabling better resolution of coastlines and straits than
most existing Eocene models (∼ 3∘ ocean resolution).
The topography (both land and ocean) uses the late Eocene (38 Ma)
reconstruction of Baatsen et al. (2016). This topography
is distinct from previous reconstructions (e.g. Markwick,
2007) in that it uses a palaeomagnetic reference frame to position the
continents (Torsvik et al., 2012; van Hinsbergen
et al., 2015), rather than a hotspot reference frame
(Seton et al., 2012). Continental elevation is based on
ETOPO modern-day topography, which is then relocated to its 38 Ma position
by plate tectonic motion using GPlates. For the deep ocean, an age–depth
relationship is applied (Müller et al., 2008) and adjusted
to the palaeomagnetic reference frame. Manual adjustments are then applied to
areas where elevation changes are well constrained by geological evidence.
Specific regions of adjustment include Antarctica, the Himalayas, the
Amazon, Turgai Strait and the Tethys Sea. Where palaeo-elevation data is
missing or unknown, the topography defaults to modern-day elevation. One
region of uncertainty is in the gateway between the Arctic and Atlantic
Ocean, where other reconstructions have a more constricted throughflow (e.g.
Fig. 41 of Markwick, 2007). In particular, deepening of the
Greenland–Scotland Ridge is hypothesised as a trigger of North Atlantic deep
water formation at the EOT, via its impact on freshwater transport from the
Arctic into the Atlantic (Abelson and Erez, 2017;
Stärz et al., 2017).
Manual adjustments were made to the topography to ensure that all straits
are at least two grid cells wide, so that they have corresponding velocity
grid points, and no ocean grid cell is isolated. The resulting ocean
bathymetry is shown in Fig. 1a. In the
atmosphere, the topography is interpolated onto the horizontal grid and then
smoothed using a 3-point mean filter to ensure a smoother interaction with
the wind field. This filtering was mainly needed on the Antarctic continent,
due to convergence of meridians on the topography grid, which caused
numerical noise in the wind field during initial testing. This filter will
likely be removed outside of the polar regions in future versions of the
model. The resulting topography is shown in Fig. 1b. Vegetation types are founded on a dataset based on
Sewall et al. (2000), with modifications where data are
available from Thorn and DeConto (2006); Utescher and Mosbrugger (2007); and Gomes Rodrigues et al. (2012).
The Sewall et al. (2000) dataset was originally
configured for the CESM plant functional types, and we have adapted these to
the corresponding vegetation types in CM2.1. The river run-off is determined
by a relocation map, where each land point is assigned a corresponding
coastal location for returning run-off to the ocean. This relocation map was
determined by a downslope relocation algorithm from the model topography.
Aerosol forcing was taken from the Eocene (55 Ma) reconstruction of
Herold et al. (2014), and adapted to our model's input
types.
Vertical mixing scheme
We use a simplified version of the tidal mixing scheme of
Simmons et al. (2004). The CM2.1 code distribution
provides a heterogeneous distribution of seafloor roughness amplitude based
on high resolution maps of the modern seafloor. We set this to be uniform
for simplicity, since estimating the late Eocene seafloor roughness is not
straightforward. We thus calculate an average roughness amplitude of 210.512 (the h parameter from Eq. (1) of Simmons et al. 2004),
and set a background
diffusivity of 1.0 × 10-5 m2 s-1. In parallel we also
simulate the model using the standard Bryan–Lewis diffusivity in CM2.1,
commonly used in deep-time palaeoclimate studies (e.g.
Lunt et al., 2012). We prefer a bottom roughness mixing scheme to
Bryan–Lewis, since the former is more physically realistic. The Bryan–Lewis
scheme uses the following values:
Keqz=10-4[0.65+1.15πtan-1(4.5×10-3z-2500)]Kpolez=10-4[0.75+0.95πtan-1(4.5×10-3z-2500)],
where Keq and Kpole are the equatorial and polar diffusivities
respectively, in m2 s-1. The depth z is in metres, which has a transition
level at 2500 m as shown above. The low latitudes use Keq and the high
latitudes use Kpole, with a transition at 35∘ latitude.
Initial conditions and spin-up
The initial conditions of the ocean model are configured to a modified
version of the DeepMIP experimental design (Lunt et al.,
2017). The temperature and salinity are set to the following profiles:
T[∘C]=(5000-z)/5000*20cos(φ)+10;ifz≤5000m.=10;ifz>5000m.S[psu]=35.0,
where φ is latitude, z is depth of the ocean (metres below surface
– positive downwards).
The model is run at three different levels of CO2: 400, 800 and 1600 ppm, where 800 ppm is deemed as the control simulation. The atmosphere and
land surface components are initialised from a previous test simulation.
We used the following procedure to spin up the model simulations:
The model is run in coupled mode for 50 years.
The last 10 years of this coupled simulation are used to generate surface
boundary conditions for an ocean and sea ice only simulation.
The model is run in ocean and sea ice only model for 500 years, using a
10-year repeating pattern of forcing based on the CORE protocol
(Griffies et al., 2009). During this simulation,
we doubled the ocean tracer time step with respect to the momentum time
step.
The resulting ocean state is used to reinitialise step 1.
After six cycles of the above (i.e. a total of 3300 model years), the
simulation was continued in fully coupled mode (with no acceleration of
tracers) for a further 3200 years.
In order to assess the robustness of this acceleration method, we also ran
the control simulation from the same initial conditions for 1000 years in
coupled mode only. This coupled-only simulation yielded an ocean state that
was very similar in key metrics (i.e. temperature, salinity, age tracer) to
the iteratively coupled ocean run at year 2000 (see
Fig. 2). In other words, the iterative coupling
procedure achieved a climate state that was similar to that of an ordinary
coupled run, with roughly twice the number of model years needed to reach
the same state. This approximately cancels out the computational speed-up
achieved by the decoupling procedure. We therefore abandoned the iterative
coupling procedure after six cycles and completed the spin-up in coupled mode.
Spin-up evolution of the 400, 800 and 1600 ppm
simulations, showing the global mean temperature evolution at (a) the
surface, (b) 2000 m, (c) 4000 m and (d) the time rate of change at 4000 m.
Panels (a–c) also include the 1000-year coupled-only run at 800 ppm,
illustrating the faster evolution achieved in coupled mode. (e) Meridional
overturning circulation (MOC) indices for each simulation: positive values
indicate the maximum overturning value in the Northern Hemisphere (NH), while
negative values indicate the maximum overturning value in the Southern
Hemisphere (SH). (f) Arctic salinity evolution, showing the total (top three
curves) and the surface average (bottom three curves). Panels (d) and (e) are
filtered using a 21-year running mean.
The spin-up evolution of ocean temperature across the three levels of
CO2 is shown in Fig. 2. The simulations all
have surface climates in quasi-equilibrium, although the deep ocean is
gradually cooling in all cases. All simulations also have a temperature trend of
less than 0.1 ∘C per century at 4000 m, although the warmer climates
are trending more slowly than the colder ones. This is because the initial
conditions are designed to be warmer rather than colder than equilibrium, so
that convection can readily occur due to surface cooling. The step changes
in SST over the first 3300 years are due to the coupling/decoupling
procedure described above. We also examined the evolution of the meridional
overturning circulation during spin-up (Fig. 2e).
This shows that the magnitude of Northern Hemisphere overturning remains
steady over the last ∼ 2000 years, while the Southern
Hemisphere overturning is gradually reducing in the 1600 and 400 ppm
cases, and is steady in the 800 ppm case. Southern Ocean sinking is
discussed further in Sects. 3 and 4. The trend in Arctic salinity is also
shown in Fig. 2f. This indicates a rapid
adjustment towards a fresh Arctic surface in all three cases, followed by
minimal change over the last 3000 years.
Net radiative forcing at the top of atmosphere (TOA)
plotted against surface air temperature. Data points are derived from each
10-year average during the iterative coupling phase of the spin-up, and then
from each 100-year average in the fully coupled phase. The first 10 years
are omitted in all three cases since the net radiative forcing is
temporarily positive. The time evolution goes from the bottom right to the top left
in all cases (i.e. a cooling trend).
Control climate state at 800 ppm, showing (a) sea
surface temperature (SST), (b) surface air temperature (SAT), (c) sea
surface salinity (SSS), (d) evaporation minus precipitation, (e) barotropic
streamfunction and (f) zonal wind stress.
(a) First empirical orthogonal
function (EOF1) of the monthly sea surface temperature anomaly (SST), and
(b) precipitation. From the EOF1 of SST we then define an El Niño index
region from 170∘ E to 140∘ W, and from 5∘ S to
5∘ N, as shown in the green box in (a). This region was chosen as
an “El Niño index” to be representative of strongest variability in
Pacific SST. (c) El Niño index variability in monthly SST anomaly.
(d) Annual mean east–west Pacific SST difference.
One way to estimate the final equilibrium in a transient climate simulation
is to derive a “Gregory plot” (Gregory et al., 2004),
which compares the top-of-atmosphere (TOA) net radiative forcing with the
temperature anomaly. This is shown for each CO2 level in
Fig. 3. The data points are derived from 10-year
averages during the iterative coupling phase of the spin-up, and then from
100-year averages during the fully coupled phase. There is a trend towards
zero net radiative forcing during much of the spin-up, but the curves begin
to flatten before they reach zero. This makes extrapolation towards an
equilibrium value difficult, since there is no definitive slope from which
to project. In addition, we note the CM2.1 model has been documented to have
a radiative imbalance under pre-industrial conditions on the order of 0.5 W m-2 (Delworth et al., 2006). This is due to a number
of factors, including a systematic loss of energy (∼ 0.3 W m-2) in the atmosphere model, and a heat sink that occurs when water
is returned to the ocean from precipitation or run-off (∼ 0.14 W m-2). Therefore, we do not expect the model to reach a perfect
radiative forcing balance at equilibrium.
March (NH maximum) sea ice thickness of the 800 ppm run (a), and the 400 ppm run (b),
and September (SH maximum) sea ice
thickness for the 800 ppm run (c) and the 400 ppm run (d). The 1600 ppm run
is sea ice free all year.
Climate and ocean circulation
The late Eocene palaeogeography incorporates changes that affect important
features of the ocean circulation compared with the present day. The
Southern Ocean gateways of Drake Passage and the Tasman Seaway are narrow
and shallow (sill depths of around 500 m), and permit a weak eastward
circumpolar flow. Panama Seaway is open, and allows a shallow connection of
surface water between the equatorial Atlantic and Pacific oceans. The Bering
Strait is closed, so that the North Pacific has no direct connection with
the Arctic Ocean. The Arctic Ocean is connected to the North Atlantic by
narrow and shallow seaways (up to 200 m depth), which allow export of
freshwater from the Arctic into the Atlantic. Transport between the Arctic
and Atlantic is bidirectional, with surface waters travelling predominantly
southwards, and waters below 100 m travelling predominantly northwards. The
Tethys Ocean has connections to the Indian and the Atlantic, and a very
narrow and shallow connection via Turgai Strait to the Arctic (25 m depth),
allowing unidirectional southward transport of surface waters. Apart from
these gateway changes, the Pacific is wider and more open through the
Indonesian archipelago, while Australia is further south. The Atlantic is
narrower especially in the Northern Hemisphere, and in this region the
subpolar gyre is particularly constricted. These features all contribute to
a vastly different ocean circulation than present day, and we explore the
major changes below.
Surface climate
In this section, we describe the control climate (800 ppm), key features of
the circulation and how they compare with proxy evidence.
Fig. 4a and b
show the sea surface temperature (SST) and surface air temperature (SAT)
respectively. Equatorial SSTs reach as high as 38 ∘C in the
western Pacific warm pool, with tropical temperatures around
30–35 ∘C elsewhere. The warm pool extends west into the Indian
Ocean due to an open Indonesian Seaway. The meridional temperature gradient
is a little lower than in the modern climate, with the high latitude SSTs of
the North Pacific and Southern Ocean being around 15–20 ∘C, even
along the coast of Antarctica.
The sea surface salinity (Fig. 4c) indicates very
fresh conditions in the Arctic Ocean, around 20 psu. The fresh Arctic
surface waters primarily flow out into the North Atlantic, where the surface
salinities around Greenland and above ∼ 45∘ N are
generally below 30 psu. Under these conditions deep sinking does not occur
in the North Atlantic. Conversely, the high latitude North Pacific
surface water is mostly around 35 psu, apart from a limited region in the
north-east where there is strong run-off entering the ocean from the North
American mountain ranges. Under modern conditions, orography preferentially
freshens the surface water in the North Pacific over the North Atlantic
(Sinha et al., 2012; Wills and Schneider, 2015; Maffre et al., 2018). This is primarily due to its
influence on precipitation and drainage pathways, especially in the North
American mountain ranges. Notably, the low surface salinities prevent deep
water formation in the present-day North Pacific (Warren, 1983).
However, our control simulation shows a higher river run-off (0.28 Sv) going
into the Arctic than the North Atlantic (0.14 Sv) and North Pacific (0.15 Sv) (here the cut-off latitude is chosen to be 45∘ N). The same
comparison with a simulation of CM2.1 in its modern configuration produces a
river run-off of 0.10 Sv into the Arctic, only slightly higher than the North
Atlantic value of 0.07 Sv and the North Pacific value of 0.08 Sv. This high Arctic run-off,
combined with enhanced precipitation minus evaporation (P-E) forcing of the warm climate creates the very
fresh Arctic surface waters. Transport of this fresh Arctic water mass into
the North Atlantic causes a state reminiscent of the present-day North
Pacific with low surface salinity, a strong halocline and no deep water
formation.
Zonal wind stress (Fig. 4e) in the Southern Ocean
has a more zonal structure than in the present day, with a weaker standing
wave meander and a lower magnitude overall. The wind stress peak sits around
50∘ S, similar to present day, and the wind stress curl drives
subtropical gyre circulations in the South Pacific and Indian oceans
analogous to the present day, as shown in the barotropic streamfunction of
Fig. 4f. The Southern Ocean gateways – Drake
Passage and the Tasman Seaway – are open, but they are narrow and shallow
and thus allow only a weak circumpolar flow around Antarctica of 15 Sv.
The mean state of the model has an east–west temperature contrast across the
equatorial Pacific in excess of 6 ∘C, slightly larger than in the
present day. Empirical orthogonal function (EOF) analysis of global SST
indicates the strongest variability in the central equatorial Pacific
(Fig. 5a), while precipitation varies most
strongly in a dipole spanning the western equatorial Pacific and Indian
oceans (Fig. 5b). Using the first EOF of SST, we
define an “El Niño Index” region between 170∘ E and
140∘ W, and 5∘ S to 5∘ N, as being
representative of highest Pacific equatorial SST variability. This index is
analogous to the present-day Niño-3.4 index, designed to bracket the
region of highest El Niño–Southern Oscillation (ENSO) variability. This
index indicates the occurrence of multi-year El Niño and La Niña
events (Fig. 5c), defined by a threshold of
+0.5 ∘C for an El Niño and -0.5 ∘C for a La
Niña, as used in present day. The magnitude of temperature deviations is
generally weaker than present-day Niño-3.4 variability. The east–west
SST gradient in the equatorial Pacific is closely anti-correlated with the
El Niño index, with a reduction in east–west gradient during El Niño
and vice versa for La Niña (Fig. 5d).
However, the magnitude of these variations is also weaker than present day.
The areal extent of the western Pacific warm pool in our model is larger
than the present day. We suggest that this partly due to fewer land barriers
in the western Pacific and a wider basin. We argue that this creates a
larger thermal inertia in the western Pacific, implying a reduction in
variance in the east to west thermal gradient. This result agrees broadly
with von der Heydt and Dijkstra (2006), who also found a
stronger western Pacific warm pool in an Oligocene model simulation.
The “Weddell Sea” equivalent in the EOT model, where deep water forms in
the model, has annual mean sea surface temperature below 10 ∘C and
develops a small area of seasonal sea ice close to the coast. The Arctic
Ocean temperatures are typically around 4–6 ∘C, and it is mostly
ice free all year round, apart from some embayments around the Siberian
coast. The average maximum monthly sea ice thickness for each hemisphere is
plotted in Fig. 6, showing March thickness for the Northern Hemisphere and
September thickness for the Southern Hemisphere. We note that there is
evidence of ice-rafted debris in the Arctic from the middle Eocene onwards
(Stickley et al., 2009), indicating the likely presence of seasonal sea ice
at this time. Our 400 ppm simulation has Arctic-wide sea ice in summer
(Fig. 6b), in agreement with this proxy evidence. However, the ice-rafted
debris emanates from the central Arctic (Stickley et al., 2009), where there
is still some sea ice in the 800 ppm case, making it difficult to
distinguish which case is more realistic.
Meridional overturning circulation
The meridional overturning circulation (Fig. 7)
shows a structure of sinking in the North Pacific and Southern Ocean. There
is no deep overturning cell in the North Atlantic, due to the surface waters
being far too fresh to allow for local deep water formation. The structure
of the Pacific overturning is analogous to modern-day North Atlantic deep
water (NADW) formation: the northern component water is warmer and saltier than the
southern component water and is heavier at the surface. This difference in
deep water properties is caused by colder winter temperatures in the
Southern Ocean sinking regions than the North Pacific, since deep water
forms exclusively in winter. The colder winter temperatures are due to the
asymmetry of land masses between the poles. One factor is that the Southern
Ocean extends further poleward than the North Pacific (in both Pacific and
Atlantic sectors). Furthermore, the Antarctic continental interior becomes
colder in winter than the Arctic Ocean, and therefore the seasonal cycle in
high latitude (> 60∘ ) surface air temperature is
stronger in the Southern Hemisphere. The Northern Hemisphere also has very
cold winters in continental interiors, but its polar winters are moderated
by the presence of an ocean and its seasonal cycle is milder.
(a) Global meridional overturning circulation (MOC),
with the Northern Hemisphere MOC split into (b) Pacific Basin and
(c) Atlantic Basin.
(a) Age tracer at 2000 m depth and (b) mixed layer
depth. Note that the age scale is saturated in the Arctic, where the strong
halocline prevents ventilation and ages are greater than 3000 years.
As in modern climate, thermobaric effects control the layering of the
abyssal ocean (Nycander et al., 2015). The Southern Hemisphere
sourced deep waters are colder and therefore more compressible than the
warmer and slightly more saline ones sourced from the Northern Hemisphere.
Thus bottom waters are predominantly formed in the Southern Ocean. Sinking
occurs in both the Pacific and Atlantic sectors of the Southern Ocean, as
indicated by the age tracer at 2000 m and winter mixed layer depths
(Fig. 8). The shallow gateways of Drake Passage
and Tasman Seaway inhibit the exchange of deep water between the basins,
preventing Pacific- and Atlantic-sourced bottom waters from directly
reaching the opposite ocean basin. The poleward heat transport is similar to
that of the modern climate (not shown), with the partition between ocean and atmosphere heat
transport being dominated by the ocean in the tropics only, and by the
atmosphere in the mid to high latitudes. As in the modern ocean, poleward
heat transport is skewed towards the Northern Hemisphere, with a peak value
of 2.2 PW (northward) and a minimum value of -0.9 PW in the Southern
Hemisphere. We suggest that this asymmetry is due to the analogous asymmetry
of deep water formation; i.e. the northern overturning cell traverses a
larger temperature difference and hence enhances northward ocean heat
transport relative to weaker temperature gradient across the southern
overturning cell. In the tropics, we do not find any warm salt-driven deep
water formation, in line with previous studies (Table 1).
Stratification-weighted diffusivity (see text)
computed over (a) global, (b) Indo-Pacific and (c) Atlantic domains for the
control (800 ppm) and Bryan–Lewis diffusivity experiments.
We also compare the results of our control run with the widely used
Bryan–Lewis mixing scheme, as used in the default case of CM2.1
(Delworth et al., 2006). The Bryan–Lewis simulation was
spun up using the same procedure as the control run, as described in Sect. 2.3. Prior to this study, we expected that changes to the bathymetry would
alter the deep ocean mixing distribution due to intensified mixing over
ridges and seamounts. Furthermore, we anticipated that changes to surface
radiative forcing would alter the thermocline structure. For these reasons
we expected that a bottom-enhanced vertical mixing scheme would be more
adaptable to late Eocene climate conditions and perform better than the
fixed vertical mixing structure of the Bryan–Lewis scheme. However, we find
that the meridional overturning circulation and stratification are largely
insensitive to this change in vertical mixing. The magnitude of the
overturning, distribution of sinking regions and age tracer were all very
similar between the mixing schemes (not shown). This was a surprising
result, given theoretical predictions of change due to bottom-enhanced
mixing (de Boer and Hogg, 2014). We do find that the abyssal
ocean is slightly cooler in the bottom-enhanced mixing scheme
(0.5 ∘C), suggesting a weaker diffusive heat penetration into the
abyss, in line with theoretical predictions. We also find abyssal salinity
differences on the order of 0.1 psu. Given the very long spin-up time of the
simulations, the lack of divergence between the schemes is surprising. In
the Arctic Ocean, the differences in deep ocean temperature and salinity are
greater. This is because the deep waters in the Arctic are very stably
stratified by the halocline and can only communicate with the surface
diffusively.
In order to examine the effective diffusivity Keff of each mixing scheme, we
computed the stratification-weighted diffusivity
(de Lavergne et al., 2016) between the two
schemes as follows:
Keff=∬KvN2dxdy∬N2dxdy
where Kv is the total vertical diffusivity, and N2 is the buoyancy
frequency. This was computed offline using monthly output of Kv and
N2 and averaged across the last 50 years of simulation (Fig. 9).
This calculation shows that the effective
mixing across the thermocline is indeed similar between the two schemes.
Mixing is slightly stronger in the Bryan–Lewis scheme between 2500 and 3000 m, while mixing near the seafloor is indeed stronger in the bottom-enhanced
scheme. There is also larger effective mixing in the Atlantic than the
Indo-Pacific in both schemes, and also a larger difference between the
bottom-enhanced mixing and the Bryan–Lewis scheme in the Atlantic.
We further note that the Simmons et al. (2004) mixing
scheme used in our control run assumes that the energy input to the mixing
within each grid column is proportional to the simulated bottom
stratification. An increase in bottom stratification (e.g. as a result of
increased surface density gradients) is, therefore, immediately paralleled by an
increase in mixing energy. This can have the effect of maintaining
approximately constant diffusivities and circulation rates despite larger
density contrasts. In this way, the Simmons et al. (2004) scheme is similar to the Bryan–Lewis scheme in that it fixes
diffusivity rather than mixing energy. A better approach may be to fix the
energy consumed by vertical mixing, since energy constraints provide a
physically consistent framework to relate the drivers of mixing to the
abyssal circulation (de Lavergne et al., 2016).
Comparison of model 1600 ppm (red), 800 ppm (orange)
and 400 ppm (green) simulations with a compilation of SST proxy data (black
circles) from 38 to 34 Ma (Baatsen et al., 2018a, Table A2). The compilation
combines UK'37, TEX86, Mg / Ca, δ18O and Δ47 data – see text for
further references. The model temperature for each proxy location is shown
by the coloured circles. The zonal mean and standard deviation are shown by
the solid lines and shaded areas respectively. Also shown in blue is an
equivalent distribution of the zonal mean and standard deviation from the modern
World Ocean Atlas (Locarnini et al., 2013).
Sea surface temperature (SST) difference between (a) 1600–800 ppm and (b) 800–400 ppm and surface air temperature (SAT)
difference between (c) 1600–800 ppm and (d) 800–400 ppm.
Surface salinity changes between (a) 1600–800 ppm
and (b) 800–400 ppm, with the corresponding changes in evaporation minus
precipitation for (c) 1600–800 ppm and (d) 800–400 ppm. Note the
nonlinear scale in panels (c–d).
Comparison with proxy data
A comparison of proxy data with the model SSTs for the 400, 800 and 1600 ppm
experiments is shown in Fig. 10. The data are
taken from the SST proxy compilation from 38 to 34 Ma shown in Table A2 of
Baatsen et al. (2018a). The compilation combines
UK'37 , TEX86, Mg / Ca, δ18O and Δ47 data
(Kamp et al., 1990; Pearson et al., 2001, 2007; Tripati and Zachos, 2002; Kobashi et al., 2004; Bijl et al., 2009; Liu et al., 2009; Okafor et al., 2009; Wade et al., 2012; Douglas et al., 2014; Petersen and Schrag,
2015; Hines et al., 2017; Evans et al., 2018). The palaeolocations of these proxies are
consistent with the model boundary conditions (Baatsen et al.,
2016). Figure 10 shows the proxy data points
compared with the nearest model ocean point, and the model zonal mean
temperature. The 1600 and 800 ppm cases are too warm compared with the
proxies in the tropics. In the control case (800 ppm) the model is roughly
3 ∘C warmer at the equator. However, the TEX86 data of Pearson et al. (2007) from
Tanzania (∼ 32 ∘C) agree better with the
control case. In the mid to high latitudes, the control case is colder than
the proxy estimates. The 1600 ppm case agrees better with the mid to high latitude
proxy estimates, however its warm bias at the equator is very
large. In the 400 ppm simulation, the tropical temperatures agree well with
the proxies, while in the midlatitudes, it is too cold.
The model's inability to capture the low meridional temperature gradient in
the proxies is a common problem in simulating warm climates of the Eocene
(Huber and Caballero, 2011). However, several aspects of the
proxy data could help to improve this situation. First, alternative
calibrations of the TEX86 data from Tanzania may yield temperatures as
high as 35 ∘C in the late Eocene (Bijl et al.,
2009), which would reduce the low gradient problem. We also note that there
are no available proxy estimates from 38 to 34 Ma in the western Pacific
warm pool, where the ocean's warmest temperatures are found. Middle Eocene
(42 to 38 Ma) SST estimates from Java are as high as 35 to 36 ∘C
(Evans et al., 2018). Second, high
latitude TEX86 data may be affected by summer bias due to seasonal
growth of their underlying species (Kim et al., 2010),
which may have the effect of reducing the meridional temperature gradient
seen in the proxies.
Climate sensitivity to CO2Temperature response
The global mean SST in the control run (800 ppm) is 27.8 ∘C;
halving CO2 to 400 ppm produces 3.5 ∘C of cooling, and
doubling CO2 to 1600 ppm yields 4.2 ∘C of warming. The global
mean SAT in the control is 25.6 ∘C; halving CO2 to 400 ppm
produces 4.0 ∘C of cooling, and doubling produces 4.8 ∘C
of warming. The higher sensitivity of temperature change in the atmosphere
is due to both stronger polar amplification in the atmosphere than the
ocean, and a stronger temperature change over land than over the ocean. In
its modern configuration, CM2.1 has an equilibrium sensitivity of
3.4 ∘C (Winton et al., 2010),
although we note that this sensitivity may be different when using the present
lower resolution atmospheric-model component. The lower sensitivity in the
modern case may be due to many factors, for example higher albedo due to the
presence of ice sheets, and the higher percentage of land versus ocean. The
net incoming shortwave radiation from our late Eocene model has an extra 12 W m-2 of insolation compared with the modern CM2.1, indicating a
significantly lower albedo. Furthermore, our simulations indicate some state
dependence, with higher sensitivity at higher CO2.
The response of SST and SAT to CO2 changes is shown in
Fig. 11. In the 1600 ppm case
(Fig. 11a, c), the tropics warm by between
3 and 4 ∘C, with a strong amplification in both of the polar regions.
SAT over the Arctic Ocean and over Antarctica are 7–9 ∘C warmer,
with the magnitude of polar amplification being similar in both hemispheres.
This approximate doubling of warming in the polar regions aligns with
expectations from idealised climate modelling of the polar amplification
response (Alexeev et al., 2005). The polar amplification
is a combination of the local radiative forcing and the enhanced moist
energy transport from the tropics. The Arctic Ocean warming is especially
pronounced in the 1600 ppm case, despite the absence of ice–albedo
feedbacks. Cloud radiation feedbacks may provide additional polar
amplification at high levels of CO2 (Abbot et al.,
2009). Warming over land shows a mixed response; warming over North America
and Siberia is enhanced compared with oceanic regions at the same latitude,
as is the case in southern Africa. By contrast, other regions such as
Australia, equatorial Africa and parts of South America show little change
to the land–ocean contrast. This is linked directly with the land–ocean
monsoon response in low to mid latitudes. Regions of enhanced land warming
are closely associated with enhanced drying, while regions of reduced land
warming generally also become wetter, as shown in the P-E forcing
(Fig. 12). In the higher latitudes, the monsoonal
forcing is weaker, and thus the P-E changes are not as closely linked with
the temperature response.
In the 400 ppm case (Fig. 11b, d), the magnitude
of tropical cooling and polar amplification are approximately as strong as in the
warming case. The Antarctic SAT response is, however, somewhat stronger than
the warming case due to the triggering of snow albedo feedbacks. Under 400
ppm, a far greater proportion of the Antarctic continent is covered in snow
in winter. The model does not sustain freezing temperatures over Antarctica
in summer. However, this should not be directly interpreted as evidence that
an ice sheet cannot be triggered under this climate; the model lacks the
necessary feedbacks to simulate the accumulation of long term snow and ice
(e.g. Gasson et al., 2014). One area of strong cooling
occurs in the Weddell Sea, where SSTs drop by ∼ 8 ∘C, with a commensurate change in SAT. There is also a cessation of deep
water formation in this region, and the formation of seasonal sea ice in the
400 ppm case (see Fig. 6). The Pacific sector of
the Southern Ocean cools to a similar magnitude as the warming case. The SST
cooling in the Arctic is less pronounced than in the warming case, simply
because the sea surface freezes for much of the year and therefore the
surface change is limited to ∼ 4 ∘C. The SAT
response in the Arctic is similarly strong, showing that the atmosphere
polar amplification can be much stronger than the Arctic SST change.
Salinity and hydrological cycle
In response to a doubling of CO2 to 1600 ppm, surface salinity
decreases markedly over the mid and high northern latitudes as shown in
Fig. 12a. Arctic Ocean salinities decrease by
around 1 psu, due to the enhancement of the high latitude precipitation and
river run-off into the Arctic. The river run-off accumulates from North
America, Europe and Siberia, so the net effect of the increased P-E
(Fig. 12c) on the salinity is particularly
concentrated in the Arctic. This Arctic freshening also has a marked impact
on North Atlantic salinity. Conversely when CO2 is halved, Arctic
salinities are enhanced by 1–2 psu, due to the weakening of precipitation
and run-off into the Arctic.
Salinity changes in other regions also reflect changes in the P-E balance to
a lesser extent. The north-east Pacific freshens in response to warming and
becomes saltier in response to cooling, due to the same forcing mechanisms.
The interiors of the gyre circulations in the Pacific show signatures of the
P-E forcing, but the salinity differences are not as clearly correlated
here, due to advection and lateral mixing altering the signal. One area of
notable increase in salinity under the warming scenario is the southern
tropical Pacific, where salinity is greatly enhanced. This corresponds to an
area of net drying in the P-E fields, but other regions do not show nearly
the same salinity response to the same magnitude of forcing. Clearly
advection feedbacks are at play, as also evidenced by the net freshening of
the South Atlantic in the 1600 ppm case. This occurs despite a net decrease
in P-E over the subtropical South Atlantic gyre. One possible explanation is
the enhanced African monsoon, which leads to large river run-off increases
from West Africa into the Atlantic. Furthermore, sinking in the Atlantic
sector of the Southern Ocean decreases markedly in the 1600 ppm case, which
is evident in both the mixed layer depths and age tracers at intermediate
depth (not shown). This reduction in deep water formation in turn reduces
the salt advection into the Weddell Sea, leading to freshening at the
surface. In the 400 ppm case, the tropical South Atlantic becomes saltier,
which is likely due to a decrease in river run-off from West Africa. However,
the Weddell Sea becomes fresher in this case, due to a reduction of sinking
in the South Atlantic and its associated salt advection.
Greenhouse warming enhances the hydrological cycle, due to the increased
capacity of the atmosphere to hold water vapour. This effect is known as the
Clausius–Clayperon relation. In the zonal mean, this effect has been shown
to create a pattern of change where wet regions get wetter and dry regions
get drier. A simple scaling relation can be derived that relates the change
(denoted by δ) in precipitation minus evaporation (P-E) to its
original distribution (Held and Soden, 2006):
δ(P-E)=αδT(P-E),
where α is a constant and δT is the change in temperature.
This scaling relation has been shown to hold true over ocean regions in
modern observations (Durack et al., 2012), but not
over land where the dynamics of rainfall distribution are more complex
(Greve et al., 2014). In line with the above scaling argument,
the oceanic patterns of wet vs dry regions (Fig. 4d) are enhanced in most ocean regions when CO2 is doubled. The
equatorial response of P-E has some exceptions to the scaling relation, as
the western Pacific warm pool and equatorial cold tongue do not respond
according to the same paradigm. However, we note that the equatorial Pacific
is subject to a strong zonal asymmetry due to El Niño variability, and these
patterns do not follow the scaling relation as cleanly as the latitudinal
relationship even in modern climate models (Held and
Soden, 2006).
The basin integrated forcing of P-E and river run-off upon the surface ocean
is summarised in Fig. 13. This illustrates the
combined effects of P-E forcing and river run-off in creating much fresher
Arctic conditions compared to present day. In comparing the control run (800 ppm) with the 1600 and 400 ppm cases, the warmer simulations give rise
to increased freshwater forcing in the North Pacific to a similar degree as
in the Arctic. However, the Arctic Ocean has a much smaller surface area and
restricted outflow; therefore the impacts on surface salinity of this
forcing are strongest. The connection of Arctic surface waters to the
Atlantic gives rise to large salinity changes in the North Atlantic, whereas
the North Pacific salinity is less impacted by this change. Furthermore, the
freshwater forcing is concentrated in the north-east of the basin, away from
the deep water formation region in the north-west. Thus, the North Pacific is
able to maintain deep water formation despite the enhanced freshwater
forcing.
(a) Precipitation minus evaporation into the North
Pacific and North Atlantic (both with a cut-off latitude of 45∘ N)
and the Arctic Ocean, comparing a modern CM2.1 simulation with the late
Eocene model at 400, 800 and 1600 ppm. (b) River run-off with the same
breakdown as in (a). (c) The total of precipitation minus evaporation plus run-off.
Global meridional overturning circulation (MOC) for
CO2 levels of (a) 1600 ppm, (b) 800 ppm and (c) 400 ppm.
Temperature and salinity properties in the sinking
regions, showing surface values during winter months: July to September in
the Southern Hemisphere and January to March in the Northern Hemisphere.
Values are taken south of 60∘ S in the Southern Hemisphere and
north of 50∘ N in the North Pacific. Shelf values less than 200 m
depth are excluded. Bold circles indicate regional surface means, lighter
dots indicate individual grid boxes and bold squares indicate regional
annual means at 2000 m. Contours of surface density are overlaid in each
case.
Overall, these changes in salinity are not enough to alter the preferred
regions of northern sinking. The palaeogeography employed here robustly
generates sinking in the North Pacific and in the Southern Ocean. In all
cases, the northern cell is warmer and saltier, and the southern cell is
colder and fresher and thus is heavier in the abyss, as in the present day.
The cooling-induced weakening of the hydrological cycle does bring the North
Atlantic closer to a regime of deep water formation, but these changes of
1–2 psu are not enough to trigger sinking. If we hypothesise that NADW did
form at the EOT (Borrelli et al., 2014), then other
mechanisms are needed to first raise the control state salinities in the
North Atlantic.
Meridional overturning circulation response
Meridional overturning circulation is weaker in the 1600 ppm case, by around
10 Sv in the northern cell, and up to 15 Sv in the southern cell
(Fig. 14a). Although this is a substantial
reduction, the overall structure remains similar to that seen in the control
run (Fig. 7), with sinking in the North Pacific
and Southern Ocean, albeit with a lower magnitude. The reduction in
magnitude may be partly due to the 400 and 800 ppm runs being further from
equilibrium, and therefore the deep ocean is expelling more heat. Otherwise,
the high CO2 case has a lower meridional temperature gradient, which
may alter the forcing of the inter-hemispheric MOC
(Wolfe and Cessi, 2014). In the 400 ppm case, the
overall magnitude of the MOC is similar in magnitude to the control, however
there is a northward shift of the southern cell, which manifests as a
positive change in the high southern latitudes in
Fig. 14b. As the climate cools, a stable
halocline develops in the Weddell Sea which ceases to form deep/bottom
water. In winter, however, the halocline supports the growth of seasonal sea
ice in the region.
In the control 800 ppm simulation, the surface waters in the Ross Sea are
slightly warmer and more saline than those in the Weddell Sea, but they have
densities close enough to allow both regions to induce deep convection
in winter (Fig. 15). With lowered temperatures in
400 ppm simulation, the nonlinear dependence on the temperature of sea water
density will, for the same temperature decrease, cause a larger density
increase of the warmer Ross Sea surface waters. Presumably, this feature
combined with the salt-advection feedback enhances the surface density
difference between the two regions to the point where deep convection ceases
in the Weddell Sea. As this transition occurs a stable halocline develops,
supporting the growth of seasonal sea ice in the Weddell Sea
(Goosse and Zunz, 2014). However, the net sea ice formation is
too weak in the 400 ppm simulation to produce any brine enriched deep water,
which is an important feature for the present-day dense-water formation
around Antarctica. In the 1600 ppm case, freshening occurs due to enhancement
of the hydrological cycle, leading to cessation of Weddell Sea convection.
Vigorous sinking occurs in the South Pacific in all cases, analogous to Ross
Sea convection in the present day. This area is more predisposed to sinking
due to a relative lack of river run-off ending up in that region, while the
strong subpolar gyre circulation allows lower latitude water to maintain a
salt feedback into the region. North Pacific surface waters are warmer and
saltier than in the Southern Ocean sinking regions, and thus establish a
deep water mass of comparable density to that of the Southern Ocean.
However, the Southern Ocean water mass, having lower temperature, becomes
the densest water mass in the abyss due to thermobaric effects.
Summary and conclusion
This work presents a new configuration of the GFDL CM2.1 climate model,
configured for the late Eocene. To our knowledge, the ∼ 1∘ ocean resolution is the highest ocean resolution in a coupled
model of the late Eocene to date. The lower resolution atmosphere used in
this study also
ensures computational efficiency for long timescale simulations. This
represents a significant step forward in resolution and accuracy in the
representation of late Eocene palaeogeography. However, we note that a
parallel study currently under review (Baatsen et al., 2018a)
has also developed a late Eocene (38 Ma) climate model at 1∘
resolution. Their model shows qualitatively similar climatologies to ours,
albeit with slightly higher greenhouse gas concentrations. Under this
configuration, important gateways known to impact the ocean circulation at
the EOT are better represented than they have been in the past. The model is
spun up for 6500 years from relatively warm initial conditions. This leads
to a quasi-equilibrium steady state in surface climate, while the deep ocean
is still gradually cooling at a rate of less than 0.1 ∘C per
century. The model shows relatively high sensitivity to CO2 forcing,
and a warm control climate under 800 ppm CO2. There is some agreement
with SST proxies, although the model still does not capture the very low
gradients from equator to pole implied by the available proxy records.
The model exhibits El Niño-like variability in the equatorial Pacific,
although the western Pacific warm pool is larger and more persistent than
present day. This creates a La Niña-like background state and more
robust east–west temperature difference. There are still major uncertainties
and areas for future improvement in the model. These include possible
improvements to the aerosols, vegetation, soil and river run-off schemes,
which have been configured based on an opportunistic adaptation of available
data. The sensitivity of these schemes has not been explored here, and may
provide avenues for future improvement. For example, the contribution of
land albedo, itself a function of vegetation and soil properties, may have
important impacts on both the land temperatures and the global meridional
temperature gradient.
We find that the model exhibits sinking in the North Pacific and the
Southern Ocean, under all levels of CO2. The southern water mass is
colder and fresher and dominates the abyssal ocean, while the northern deep
water mass is warmer and saltier, analogous to the present-day structure of
North Atlantic deep water overlying Antarctic bottom water. The model is
configured with a bottom-enhanced mixing scheme and a uniform background
diffusivity. Sensitivity tests indicate that using a Bryan–Lewis diffusivity
scheme, commonly used in palaeoclimate models, yields largely the same
stratification and sinking regions. The Arctic Ocean is very fresh, with
typical surface salinities of 20 psu, in agreement with Eocene salinity
proxies. The connection between this fresh water mass and the North Atlantic
prohibits the formation of North Atlantic deep water, since North Atlantic
salinities are around 25–30 psu in present-day sinking regions. These
results highlight the importance of using late Eocene palaeogeography in
modelling the EOT. Using present-day geography as a control state would not
capture the dramatic differences in salinity in the northern ocean basins
from the present day, which in turn greatly alter the preferred regions of
sinking.
In response to CO2 forcing, North Atlantic salinity clearly responds to
a strengthening of the hydrological cycle under warming, and weakening under
cooling. The net effect is a decrease of 1 psu in the North Atlantic for a
doubling of CO2, and an increase of 1–2 psu for a halving of CO2.
This effect is not enough to trigger North Atlantic sinking in the model.
However, it does suggest that if the climate were closer to a North Atlantic
sinking regime at the EOT due to factors not captured here, CO2 cooling
could provide a trigger. This model provides a platform for further
sensitivity studies in altering aspects of the palaeogeography. Changes to
the Southern Ocean gateways, the Arctic gateways and the imposition of an ice
sheet have all been shown to have significant impacts on the ocean
circulation and climate in previous modelling studies. Deepening of the
Greenland–Scotland Ridge around the EOT has been suggested as a potential
trigger of North Atlantic sinking (Abelson and Erez, 2017;
Stärz et al., 2017). It is important to investigate such changes in a
palaeoclimate model with an accurate representation of late Eocene
palaeogeography, so that changes across the EOT can be referenced to
circulation, temperature and salinity characteristics that are appropriate
to that time period.
Data availability
The data presented in this paper are available in the
Supplement.
The supplement related to this article is available online at: https://doi.org/10.5194/cp-14-789-2018-supplement.
Competing interests
The authors declare that
they have no conflict of interest.
Acknowledgements
This work was supported by the Bolin Centre for Climate Research, research
areas 1 and 6, and the Swedish Research Council, project 2016-03912. Numerical
simulations were performed using resources provided by the Swedish National
Infrastructure for Computing (SNIC) at NSC, Linköping. MB is supported
by the Netherlands Earth System Science Centre (NESSC) and the Ministry of
Education, Culture and Science (OCW), grant number 024.002.001. The authors
thank Casimir de Lavergne and an anonymous reviewer for their constructive
comments, which helped to improve the manuscript. Model data can be made
available on request.
Edited by: Alan Haywood
Reviewed by: Casimir de Lavergne and one anonymous referee
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