Introduction
Full depth open ocean deep convection has been observed only once in the
Southern Ocean, during the Weddell Polynya of 1974–1976 .
Otherwise, the mixed layer depth in the southern subpolar gyres does not
exceed 1000 m . In state-of-the-art CMIP5 (Coupled Model Intercomparison Project Phase 5) models,
Antarctic Bottom Water (AABW) is formed via open ocean deep convection
(deeper than 2000 m and up to full-depth) which occurs most winters in the
seasonally ice-covered southern subpolar gyres and
continues throughout climate change simulations . Not only
is this process relatively unrealistic, it also leads to spurious ocean
properties, variabilities and drift e.g. and
can render reanalyses useless . and
have shown that each event of open ocean deep convection
accelerates the southern subpolar gyres and hence increases the Antarctic
Circumpolar Current (ACC) transport to unrealistically high values. With the
advent of new methods to form and export AABW more realistically
e.g., it would be desirable to
eliminate unrealistic Southern Ocean deep convection.
Here we investigate the role of ocean model parameterisations in triggering
open ocean deep convection in the Southern Ocean. Observations and models
alike show that a preconditioning mechanism, reducing the static stability of
the water column, is necessary before deep convection can commence
. According to , static instability
arises because the current generation of ocean models mixes vertically too
weakly when given heterogeneous surface conditions. Buoyancy loss and mixing
imposed by sea ice concentration less than unity are averaged over the
horizontal grid; with weak mixing, the stratification becomes unstable.
Wind-induced mixing determines the amount of turbulent kinetic energy (TKE)
penetrating the mixed layer, which if underestimated can result in a
disconnection between the surface and the rest of the water column, leading
to a surface salt accumulation and subsequent loss of static stability
. More recently, and
have demonstrated the sensitivity of high latitude mixed
layers to vertical mixing. All these studies suggest that vertical mixing
needs to be increased to reduce deep convection in models.
NEMO3.4, control run, for each grid point, maximum monthly mixed layer depth
between January 1980 and December 1989. The blue box indicates the
Riiser-Larsen Sea region, which is overlaid by the magenta box that indicates
the Weddell Sea region studied in Sect. .
The ocean model NEMO3.4 Nucleus for European Modelling of the Ocean; is used here in its configuration
ORCA025 (nominal horizontal resolution of 0.25∘, 75 vertical
levels) and is coupled to the CICE sea ice model . It will
be hereafter referred to only as NEMO. Vertical mixing of momentum and
tracers is treated using a TKE scheme , with convection
being parameterised by an enhancement of the calculated vertical eddy
diffusivity and viscosity. The ocean is forced with a prescribed atmospheric
forcing from CORE2 . The ocean model configuration used in
this study exhibits deep convection in the Weddell Sea .
Here we investigate the role of vertical mixing parameterisations on Southern
Ocean deep convection with sensitivity experiments on NEMO3.4. We detail these
experiments in Sect. , along with our methods. In
Sect. , we briefly explain the mechanism triggering deep
convection in the control experiment before showing how this mechanism is
modified in the sensitivity experiments. Note that longer, more detailed
explanations of these results are given by . We also study
the impact of the changes of parameters on the ACC and North Atlantic deep convection
in Sect. . Our results are summarised and discussed, along with
their limitations, in Sect. .
Sensitivity experiments and methods
Parameters studied
The purpose of these experiments is to identify a method to reduce the
occurrence of open ocean deep convection in the NEMO Southern Ocean. The
experiments are based on similar work performed by , who
varied 13 parameters of the TKE scheme and studied their impact on the
shallow bias in the Southern Ocean summer mixed layer depth. Note that their
experiments were performed with a different ocean resolution (1∘
compared with 0.25∘ here). Their findings established which values
to use in the most recent joint NERC (Natural Research Environment Council)–Met Office configuration of the ocean
model, “GO5” , which provide the base settings used in
our experiments. Our “Control” experiment is the GO5 run “amhih” (see
Code availability at the end of this manuscript). Following the findings of
and , we increase and decrease just
three parameters as detailed below. The background diffusivity experiments
extend throughout the entire period of available CORE2 atmospheric forcing
(27 years, 1980–2006) while the other experiments have run for 10 years
(1980–1989). All the experiments are summarised in Table .
Sensitivity experiments performed on NEMO3.4, detailed in
Sect. : “Langmuir” experiments look at Langmuir turbulence
velocity scale, “Gamma” at the penetration of an additional turbulent
kinetic energy term below the mixed layer, “Knoprof” and “KProf” at
background diffusivity. “I” indicates that the parameter was increased
compared to the reference value, “D” that it was decreased. The parameters
column identifies the shorthand name used in the NEMO simulation name list.
Note that the “Control” run is 10 years long.
name
parameter
value
value in “Control”
run length
LangmuirD
ln_lc
false
true
10 yr
LangmuirI
rn_lc
0.20
0.15
10 yr
GammaD
rn_efr
0.005
0.05
10 yr
GammaI
rn_efr
0.095
0.05
10 yr
KnoProfD
rn_avt0
1.0 × 10-5
1.2 × 10-5
27 yr
KProf
nn_avb
1
0
27 yr
KProfI
nn_avb
1
0
27 yr
rn_avt0
1.3 × 10-5
1.2 × 10-5
Langmuir turbulence velocity scale “cLC”
Langmuir turbulence is represented in NEMO by the parameterization of
, which appears as an additional production term in the TKE
budget equation:
de‾dt=W3L,
where W is a velocity scale, taken to be the maximum downwelling velocity
of the Langmuir cell, and L is a length scale, taken to be the vertical
extent of the cell. A sinusoidal profile is assumed for the cell so that:
W=cV10=cLCVs|z=0sin-ΠzLfor-z≤L,W=0for-z>L,
where Vs|z=0=0.016V10 is the surface value of the Stokes
drift for a fully developed sea , V10 is the 10 m wind
speed and cLC is a scaling coefficient, suggested by
to be between 0.15 and 0.2 based on a comparison with Large
Eddy Simulation results.
Averaging over 1982–1985 and between 60 and 45∘ S in the Southern
Ocean (i.e. north of the deep convection regions), show
that increasing cLC deepens the mixed layer throughout the year,
decreasing the summer–autumn shallow bias but increasing the winter–spring
deep bias. Here we test three cases:
Langmuir turbulence parameterisation turned off (“LangmuirD”)
cLC=0.15: control experiment (“Control”)
cLC=0.20: case with increased vertical velocity scale (“LangmuirI”).
Near-inertial wave breaking TKE scaling “γ”
NEMO features an ad hoc parameterisation of the mixing due to the breaking of
near-inertial waves excited by high-frequency winds,
e¯inertial, which is added to the time-integrated TKE
:
e¯(t+Δt,z)=∫tt+Δt(∂e¯∂t(z))+e¯inertial(t,z).
e¯inertial is defined as
e¯inertial(t,z)=γe¯|z=0expz/λ,
where λ is an exponential decay scale, set as 10 m globally
in the GO5 configuration, and γ is the parameter varied here (fraction
of TKE penetrating below the mixed layer).
tested values of γ ranging from 0.005 to 0.095
(default 0.05). Considering the zonal average of the Southern Ocean between
60 and 45∘ S only, mixed layer depth biases decrease in NEMO in
summer and winter as γ increases. As for cLC, these
averages do not include the Southern Ocean deep convection areas. The value
of γ for the real ocean is unknown, hence the range of values
suggested by is purely numerical. In addition to the
control, we consider only the two most extreme values:
γ=0.005: case with decreased extra TKE mixing (“GammaD”)
γ=0.050: control experiment (“Control”)
γ=0.095: case with increased extra TKE mixing (“GammaI”).
Background diffusivity profile and surface value
Unresolved and otherwise unparameterised vertical mixing processes are
represented by a background vertical eddy diffusivity. In sensitivity
experiments on ORCA025 (i.e. the same resolution of NEMO as used here),
increased the background diffusivity (constant through
depth) from 1.0×10-5 to 1.2×10-5m2s-1 and found a significant surface
freshening and increased stratification in the Arctic. They did not detail
their results for the Southern Ocean.
In NEMO, it is also possible to change the shape of the background
diffusivity profile. Two shapes are implemented: the background diffusivity
can either be constant through depth, or increase linearly with depth
diffusivity reaches 10 times the surface value at 4000 m
depth;. Unpublished experiments with HiGEM High-resolution Global Environmental
Model;, whose ocean model is not NEMO, suggest that the open
ocean deep convection area is reduced in both southern subpolar gyres when
the background diffusivity increases linearly instead of being constant with
depth. Here we test the effects of modifying the vertical profile and/or the
surface value of the background diffusivity:
constant profile, surface diffusivity = 1.0×10-5m2s-1
(“KnoprofD”);
constant profile, surface diffusivity = 1.2×10-5m2s-1, control experiment
(“Control”);
linear profile, surface diffusivity = 1.2×10-5m2s-1
(“Kprof”);
linear profile, surface diffusivity = 1.3×10-5m2s-1 (“KprofI”).
Methods
To assess the state of the ocean, we use the potential temperature (hereafter
referred to as temperature only), salinity, sea ice concentration and mixed
layer depth (MLD) diagnostics from NEMO, in its configuration ORCA025
(horizontal resolution of 0.25∘) and with the GO5.0 default settings
. We compute the potential density (hereafter referred to as
density only) relative to the surface σθ using the equation of
state EOS80 . The MLD is determined in the model using a
density σθ threshold of 0.01 kg m-3 from
the 10 m depth value . The observed global MLD was
obtained from the climatology of . Following
, we consider that there is open ocean deep convection in
the Southern Ocean (latitude south of 50∘S) if the local monthly mean
MLD reaches at least 2000 m where the bathymetry is deeper than
3000 m. We compute the total area of deep convection as the sum of
individual model grid cell areas where the MLD exceeds 2000 m.
We define an open ocean polynya as a region where the sea ice area fraction
is locally less than 0.15, surrounded by a zone where the sea ice area
fraction is greater than 0.15 (i.e. not directly connected to the ice-free
ocean). Likewise, we compute the total polynya area as the sum of the areas
of connected model grid cells where the criteria for a polynya are met.
Anomalies are calculated relative to the first half of the simulation
(January 1980–December 1984), before the onset of deep convection.
Anomalies are considered significant if they are larger than the monthly
temporal standard deviation during 1980–1984. To account for the advection
of the anomalies by the local currents (which have little vertical shear for
the depth range studied and exhibit temporal variability, not shown), we
define the trajectory between the first two polynyas (1986 and 1987) as the
succession of monthly positions occupied by the water that was in the polynya
in September 1986, inferred from the subsequent monthly horizontal velocity
vectors.
To assess the impact of Southern Ocean deep convection on the large-scale
circulation, we calculate the Antarctic Circumpolar Current volume transport
as the annual total mean volume transport through Drake Passage. We integrate
the zonal velocity from the Antarctic Peninsula to South America, and then
over depth from the sea floor to the surface. The ACC volume transport is the
total resulting from these integrations. Following for example
, we compare the ACC volume transport with the horizontal
gradient in density across the ACC dρ, defined as the zonally averaged
density difference between 45 and 65∘ S. We also
check if the MLD in the North Atlantic is modified by our sensitivity
experiments. We consider that there is deep convection in the North Atlantic
(north of 50∘ N, 70∘ W–20∘ E) if the
local monthly mean MLD exceeds 1000 m following observations
byfor example.
Results
Mechanisms for triggering open ocean deep convection in the control run
In order to assess how our parameter changes impact the model ocean, we first
need to understand how model processes trigger Southern Ocean deep
convection. In the control run, between 1980 and 1989, the maximum MLD
exceeds 2000 m in the open ocean in the Riiser-Larsen Sea only
(Fig. , blue box), in the winter of 1987.
Chain of events leading to deep convection in 1987: (a) sea
ice concentration anomaly in June 1985, (b) MLD anomaly in September
1985, (c) and (d) sea ice concentration in September 1986
and October 1987, and (e) trajectory of the water between the two
polynyas of 1986 (dashed black line) and 1987 (continuous black line, see
text). Stippling on (a) and (b) indicates areas where the
anomalies are not significant. Dashed contours on (a) and
(b) indicate the location of the 1986 polynya. Black contours on
(a)–(d) indicate the area where the maximum MLD from
Fig. exceeds 2000 m, while grey contours indicate
the 3000 m isobath.
The process leading to deep convection starts 2 years earlier with a
decrease of the katabatic winds off the Antarctic coast. It results in a
positive and significant sea ice area anomaly in June 1985 close to the
coast, and a negative anomaly further offshore (probably due to a lack of sea
ice export) where the first polynya will open (more than 0.2, dashed black
contours on Fig. a). In agreement with ,
we find that the sea ice deficit is associated with a warm and salty anomaly
of the top 50 m (Fig. a, c). Meanwhile, the MLD is
anomalously shallow (Fig. b), which could be a consequence of
the decrease in winds. With a reduced mixed layer entrainment, a warm and
salty anomaly develops below the mixed layer (1.5∘C and
0.25 in September 1985; Fig. a, c). When ice-free in
summer, the warm salty waters are incorporated in the mixed layer, which is
no longer anomalously shallow (deepening from 48 m in January 1986 to
120 m in May). The surface waters become anomalously warm
(Fig. a), remaining above the freezing point until June 1986
(Fig. b). The warm surface waters impede sea ice formation,
resulting in the development of an open ocean polynya from August to
October 1986 (maximum extent of 22 000 km2 in September 1986;
Fig. c), north of the 1987 polynya. These surface waters are
also strongly anomalously salty and hence dense from June 1986
(Fig. c, e) and become even more so once the polynya is
open. In agreement with observations e.g., the model
polynya allows for the formation of dense water at the surface
(Fig. e, f) due to brine rejection, further destabilising
the water column and inducing an increased salinity-driven convection that
reaches a maximum depth of 827 m in September 1986 at the centre of the
polynya.
Hovmoller diagrams, as a function of depth and time, of properties
averaged over the area of the 1986 polynya (dashed black contour on
Fig. e) until August 1986, along the trajectory of
Fig. e from September 1986 to June 1987, and over the area of
the 1987 polynya (plain black contour on Fig. e) from July
1987: (a) potential temperature anomalies (relative to January 1980–December 1984), (b) potential temperature, (c) salinity
anomaly (relative to January 1980–December 1984), (d) salinity,
(e) potential density σθ anomaly (relative to
January 1980–December 1984), and (f) potential density
σθ. Grey line on (a), (c) and (e)
and white line on (b), (d) and (f) is the MLD.
The polynya and increased convection both act towards a decrease in
stratification, preconditioning the ocean for deep convection. In
October 1986, the isopycnals are vertical (Fig. f) and the
surface waters are anomalously dense (Fig. e) because of the
brine rejection in the polynya and because deep salty waters have been
brought up by convection (Fig. d). Moreover, the increased
convection reaches the layer of relatively warm waters, which enhance the
warm anomaly at the surface (from January 1987; Fig. a) so
that again the ocean surface is above freezing temperature when the sea ice
should grow (Fig. b). Meanwhile, not only are the surface
waters anomalously saline and dense because of brine rejection in winter,
they also remain anomalously saline and dense through spring and summer
(Fig. c, e), as no ice is available to be melted at the
location of the polynya. As the surface waters are anomalously warm, the
polynya reopens in July 1987 and reaches a maximum extent of
68 000 km2 in October 1987 (Fig. d). This polynya
reopens further south than the one in 1986. That is because the warm and
dense anomalies have been advected in a near-barotropic subsurface flow that
brought them to the site of the 1987 polynya (Fig. e). Also,
like real polynyas do , the 1986 model polynya and
associated increased convection had an impact over a large area, lifting
isopycnals significantly even at 66∘ S where a lower than usual
sea ice concentration can be seen (0.7 in the black contours of
Fig. c).
The conjunction of weakly stratified waters and the second polynya results in
open ocean deep convection from August to November 1987, reaching a maximum
depth of 3200 m in October 1987. In the control run, deep convection
does not restart in the subsequent years. As explained in the following
section, surface waters are not warm enough to sustain deep convection after
1987. We shall now examine how the process triggering deep convection is
modified in the parameter sensitivity experiments.
Modifications induced by the experiments
All sensitivity experiments start from the same initial conditions as the
control simulation. Considering the mean 1980–1984 sea surface temperature,
over the area that will convect (black contours on Fig. a),
from November to March the decreased parameter experiments (reduced vertical
mixing) are warmer than the control. The increased parameter experiments
(increased mixing) in contrast are colder than the control. The largest
difference from the control is found in February, over the region of deep
convection, for the TKE experiments: on average GammaD is warmer than the
control by 0.34∘C while GammaI is colder by
0.22∘C. There is no significant difference in the winter
sea surface temperature (SST), as the area is ice-covered and SST is at the
freezing point during 1980–1984 in all experiments. The winter mixed layer
is modified by the Langmuir turbulence experiments (5 m shallower
than the control in LangmuirI, 5 m deeper in LangmuirD), but not
significantly by the other experiments. This is consistent with the results
of : increasing the Langmuir turbulence velocity scale
deepened their winter MLD by 5–10 m, whereas Gamma had little
effect on it.
In the winter of 1985, all experiments show the same sea ice anomaly, at the same
location, constrained by the atmospheric forcing. The temperature anomaly of
the summer of 1986 differs significantly from that of the control run only for the
Langmuir experiments (anomaly in LangmuirD warmer by 0.3∘C
than in the control and 0.1∘C colder in LangmuirI) and for
KnoprofD (0.3∘C warmer than the control). The results of
the Langmuir experiments are consistent with those of ;
the results of the background diffusivity experiments are more surprising as
it is usually thought that the background diffusivity would not have any
effect on such shallow MLD e.g.. The regional
climatologies of the decreased parameter experiments are warmer than the
control in the summer of 1986. Their vertical mixing is weaker than the
control; hence, they accumulate warm and saline anomalies at the surface.
Across-run significant relationships between the steps leading to
the deep convection event of the winter of 1987. (a) Sea surface
temperature in June 1986 and polynya area in September 1986; (b)
polynya area and maximum depth of the mixed layer, both in September 1986;
(c) maximum depth of the mixed layer in September 1986 and sea
surface temperature in February 1987; (d) sea surface temperature in
June 1987 and polynya area in October 1987; (e) polynya area and
maximum depth of the mixed layer, both in October 1987; (f) polynya
area and deep convection area, both in October 1987. Horizontal bars on
(a) and (d) and vertical bars on (c) indicate the
standard deviation relative to spatial variability.
The warmer the ocean is in the summer and autumn of 1986, the less sea ice can
form and thus the larger the polynya in the winter of 1986 (Fig. a). The
larger the polynya, the deeper and more extensive the convection of 1986
(Fig. b). The deeper the convection, the warmer the
surface of the ocean after the convection (Fig. c), and
consequently the larger the polynya in the winter of 1987
(Fig. d). The larger the 1987 polynya, the deeper the
mixed layer and the larger the area of deep convection in 1987
(Fig. e and f). All relationships are relatively linear
and significant, albeit with some uncertainty in the surface properties. Note
that the whole process is enhanced by salinity anomalies, in particular the
larger the polynya, the saltier the surface waters and the more destabilised
the water column .
After 1987, the LangmuirD, GammaD, KnoprofD and Kprof keep convecting. This
is not surprising: the positive feedback or “temperature loop” found above,
sustained by warm water advection and brine rejection, has been responsible
for decades of deep convection every winter in other models
e.g.. As in other modelling studies
e.g., the centre of the deep convection area is
advected to the middle of the Weddell Gyre (Fig. ). There
is a very strong and significant across-run relationship between the deep
convection area of 1987 in the Riiser-Larsen Sea and the area of the 1988
polynya and deep convection (correlation of 0.90). In fact, advected by a
steady westward slope current, the warm anomalies induced by the 1987 deep
convection in the Riiser-Larsen Sea propagate to the centre of the Weddell
Gyre where deep convection occurs in 1988 and 1989 (Figs. , a). The three increased parameter experiments (and the
control) are not warm enough for their anomaly to trigger the opening of a
polynya in 1988 (i.e. they do not have sufficient heat content to prevent sea
ice formation), hence they stop convecting and their convection area is
restricted to the Riiser-Larsen Sea (Fig. ).
For each experiment, contours indicating where the maximum
1980–1989 MLD exceeds 2000 m in the Weddell Gyre region (magenta box
in Fig. ). Thin grey contours indicate the 3000 m
isobath. Note that the contours for the increased-parameter experiments and
the control mostly coincide.
ACC and North Atlantic deep convection
We have focused on the impact of changes to the vertical mixing parameters on
a specific ocean characteristic, the MLD. If these changes are to be
considered part of the solution to spurious Southern Ocean deep convection,
we must ensure that they have no adverse impact elsewhere in the global model
ocean. We found that the ACC volume transport increases following deep
convection in the Weddell Sea. Such a response had been hypothesised by
. In an 18-year run of the model BRIOS2 (Bremerhaven Regional Ice Ocean Simulations) with a
specified atmospheric forcing, they found that Weddell Sea deep convection
intensified the Weddell Gyre circulation and suggested that it would then
strengthen the ACC (theirs was prescribed). In agreement with other modelling
studies e.g., deep convection
in the Weddell Sea is associated with an increase in the horizontal density
gradient, a key driver of the ACC (Fig. ). We found a
significant correlation between the ACC volume transport and the area of deep
convection (correlation of more than 0.6 for the three 27-year
simulations, Fig. ). For each year during which deep
convection occurs, the ACC transport increases by 2–4 Sv.
27-year time series of the annual maximum ACC (black), annual mean
horizontal gradient in density dρ (green) and annual maximum area of
deep convection in the Southern Ocean (red).
We found an increase in the ACC tranport of nearly 25 Sv in the
27-year simulations. This result is consistent with the increase of
more than 20 Sv found by and the increase of
20 % found by while modelling the Weddell Polynya
in GFDL-MOM4 (Geophysical Fluid Dynamics Laboratory-Modular Ocean Model). Such an increase is an issue for the model, especially as in
our experiments the atmospheric forcing is prescribed and does not react to
deep convection. Even with a forced model, we obtain an unrealistically
strong ACC: at the end of the run, all three long simulations have exceeded
the observational range of 134–164 Sv . These
perturbed parameter experiments show that Southern Ocean deep convection
disturbs the large-scale oceanic circulation.
While deep convection in the Southern Ocean is a rather unrealistic process
occurring in models, deep convection in the North Atlantic is a key driver of
the global ocean circulation e.g.. Ideally, we would
like to minimalise Southern Ocean deep convection whilst maintaining North
Atlantic deep convection. Our results suggest that this is possible
(Fig. ). Unlike in the Southern Ocean where
a preconditioning is necessary (Fig. a), deep convection in
the North Atlantic is persistent throughout the simulation
(Fig. b), with differences between the experiments being
less than the interannual variability. The differences due to
parameterisations are not altered throughout the 27-year simulations,
consistent with the findings of ; changing the background
diffusivity has no significant effect on the northern high latitude MLD. In
general, the GammaD and GammaI simulations have respectively the largest and
smallest areas of deep convection; the other experiments show no consistency
in the North Atlantic from 1980 to 1989 (Fig. b). Our
results suggest that at least up to a decade, deep convection in the North
Atlantic is not significantly modified by the three vertical mixing
parameters found to have a large impact on the Southern Ocean deep
convection. We obtain similar results when looking at the rest of the world
ocean: outside of the deep convection regions (i.e. 60∘ S–50∘ N, not shown), the year-long area-weighted mean difference in
MLD between observations and all our simulations ranges between 4.8 m
(GammaI) and 6.5 m (GammaD), and no clear regional patterns can be
detected. That is, outside of the southern subpolar gyres, the MLD is not
significantly modified by our changes of parameters. It thus seems feasible
to reduce Southern Ocean deep convection by tuning these three parameters
without impacting the MLD elsewhere.
For each experiment, annual maximum area of deep convection in
(a) the Southern Ocean and (b) the North Atlantic.
Discussion, limitations and conclusions
We performed sensitivity experiments on a global version of NEMO3.4 with
prescribed atmospheric forcing to study the trigger of unrealistic Southern
Ocean deep convection. A complex chain of events, which preconditioned the
ocean, initiated open ocean deep convection in the Riiser-Larsen Sea in
the winter of 1987 in our simulations. It began 2 years earlier with a weakening
of katabatic winds from Antarctica resulting in a negative sea ice anomaly in
the winter of 1985 (Fig. a). These results are consistent with those
of and that highlight a strong
sensitivity of ocean models to sea ice anomalies leading to open ocean deep
convection in the Southern Ocean. Note that other ocean models have a similar
chain of events and timing, with deep convection in a small area in the
Riiser-Larsen Sea triggering a larger event in the central Weddell Sea. The
reason why the Riiser-Larsen Sea is so prone to deep convection in ocean
models remains an open question to be investigated.
Modifying the oceanic vertical mixing parameters greatly altered the extent
of Southern Ocean open ocean deep convection and the process triggering it.
The anomaly in sea ice induces anomalies in the top 50 m properties. In
agreement with , we found that the experiments with
increased mixing parameters (notably LangmuirI) resulted in smaller surface
temperature and salinity anomalies. The perturbed parameter ensemble revealed
distinct relationships between the summer and autumn sea surface temperature
and the area of the subsequent polynya in winter (Fig. ).
Once the polynya forms, the ocean enters a positive feedback loop: warmer
surface leads to larger polynyas, which have more brine rejection and hence
lead to deeper convection, which mixes up relatively warm water and leads to
even warmer surface waters. A similar process occurs in the Kiel Climate
Model , where winter deep convection persists for decades
and reoccurs on centennial timescales. In our simulations, deep convection
stops after the winter of 1987 in the control run and the increased parameter
experiments. Longer simulations would be needed to determine if the frequency
of occurrence of deep convection is generally diminished. Experiments which
reduce the vertical mixing produce strong temperature anomalies, caused by
the 1987 event, which are advected westward and result in renewed deep
convection across the centre of the Weddell Gyre in 1988
(Fig. ). These experiments exhibit convection from the
surface to the sea bed over most of the Weddell Sea in 1989 when the
simulations stop.
Our results are consistent with those of and
: deep convection increases the ACC volume transport above
the range indicated by observations (Fig. ). Whilst
increasing the model vertical mixing dramatically alters the open ocean deep
convection in the Southern Ocean, it does not significantly modify the North
Atlantic deep convection (Fig. ). This is the same
conclusion reached by , although they went to the
extent of devising a new vertical mixing scheme specifically for seasonal sea
ice regions which, after correction of a coding error, does not reduce
Southern Ocean deep convection;. Our conclusion is that a
similar result can be achieved within the parameterisation of an existing,
globally consistent, mixing scheme. Although the parameters changed here are
specific to NEMO, increasing vertical mixing in the Southern Ocean in other
models is likely to lead to reduced Southern Ocean deep convection, for the
mechanism involved would be similar to the ones highlighted in this
manuscript. Our results, however, need to be treated with some caution, as
our simulations have been too short to determine if the parameter changes
result in longer term (i.e. a spun-up model) changes to the mean state of the
global ocean simulation. Longer runs would determine the optimum balance
between a low-bias global ocean, and a Southern Ocean with reduced deep
convection. Longer simulations are also needed to obtain a more appropriate
reference period (i.e. after the model has spun up). Further observations
would also be beneficial to validate or extend the range of parameter values
used in this study, notably LangmuirI.
It is clear that the current state-of-the-art models require Southern Ocean
deep convection in order to form their Antarctic Bottom Water
. Ideally, they should also represent the actual Weddell
Polynya in hindcasts. Hence, open ocean deep convection should be reduced but
not totally suppressed. New methods, which focus on shelf overflow
parameterisations, are being developed and implemented to form Antarctic
Bottom Water in a more physically accurate manner . Some
examples are a pipe from the shelf to the open ocean ,
porous barriers and an embedded Lagrangian model
. By performing a set of vertical mixing sensitivity
experiments on the NEMO model, we have shown the general direction that
models need to take to at least reduce spurious Southern Ocean deep
convection: the vertical mixing needs to be increased, not decreased as one
might intuitively think. This paper paves the way for further model
improvement that could help ocean models to not form their bottom waters
unrealistically, by reducing open ocean deep convection in the southern
subpolar gyres.