The North Balearic Front forms the southern branch of the cyclonic
gyre in the northwestern Mediterranean Sea. Its dynamics exhibit significant seasonal variability. During autumn, the front spreads northward
during the calm wind periods and rapidly moves back southward when it is
exposed to strong northerly wind events such as the tramontane and mistral.
These strong winds considerably enhance the air–sea exchanges. To investigate
the role of air–sea exchanges in the dynamics of the North Balearic front, we
used observations and a high-resolution air–sea coupled modelling system. We
focused on a strong-wind event observed in late October 2012, which was
well-documented during the Hydrological Cycle Mediterranean Experiment. The
coupled model was able to correctly reproduce the 4 ∘C sea surface
temperature drop recorded in the frontal zone together with the observed
southwestward displacement of the front. The comparison between the weak wind
period preceding the event and the strong-wind event itself highlighted the
impact of the wind regime on the air–sea coupling. During the low-wind period
the coupling is thermal and dynamical whereas during the strong-wind period
the coupling is mainly thermal. The effect of air–sea exchanges on the
stratification variations in the frontal zone was investigated with a
stratification budget diagnosis. The stratification variations are controlled
by diabatic air–sea buoyancy flux, adiabatic Ekman buoyancy flux, and
advective processes. During the strong-wind period, the Ekman buoyancy flux
was found to be 3 times greater than the air–sea buoyancy flux and thus
played a major role in the de-stratification of the frontal zone. The role of
Ekman pumping and inertial wave in the advective processes is also discussed.
Introduction
The surface circulation in the northwestern Mediterranean Sea (NWMS) is
formed by a cyclonic oceanic gyre (Fig. ). This cyclonic gyre is
closed to the north and west by the Northern Current and
to the east by the West Corsica Current (WCC) (Fig. ). The south
branch of the surface gyre is defined by a frontal zone, the so-called North
Balearic Front (NBF). The NBF is an extension of the Balearic Current (BC),
from the Balearic Sea to the Ligurian Sea . This surface
density front (100–200 m deep) separates the warm and fresh Atlantic water (AW) which has recently entered the south of the basin from the colder
and saltier AW present in the centre of the cyclonic gyre .
This front forms a Lagrangian barrier , which plays an
important role in the nutrient budget and planktonic ecosystem
and marine ecosystem distributions
.
Schematic representation of the surface oceanic circulation in the
northwestern Mediterranean Sea. Coloured area: orography in brown and
bathymetry in blue. The directions of the prevailing mistral and tramontane
winds are indicated with black arrows. Up-front and down-front wind zones are
indicated with yellow and green ellipses, respectively.
The NBF dynamics are strongly influenced by mesoscale and sub-mesoscale
structures with significant seasonal variations . During
winter the surface front is located to the south (around 40∘ N)
because of the densification of surface waters by open-ocean deep convection
that occurs in the centre of the gyre . During summer the
light AW present in the south of the basin spreads northward and the position
of the front is further north. The short timescales variations in the
surface front can also be significant in response to the wind variability:
when the northern wind is weak, light water spreads northward over the dense
water while, when the northern wind is strong, the front shifts southward
.
This study focuses on the NBF dynamics during an autumn storm. This storm
occurs during the Intensive Observation Period (IOP) number 16
(26–29 October 2012) of the HyMeX program
. This IOP can be divided
into two parts, the first part IOP16a (26 October 2012) was dedicated to
heavy precipitation events whereas the second part IOP16b
(27–29 October 2012) focussed on strong-wind events. During IOP16a the
convergence between southwesterly and southeasterly flows initiated and
maintained strong precipitation offshore and over the southeastern French
coasts . During IOP16b a severe northerly wind event,
associated with tramontane and mistral and a large decrease in the sea surface temperature (SST) were observed at the Lion meteorological buoy
. explained this large SST drop by
the intense surface heat losses producing cooling of surface water and
vertical mixing. showed that the large decrease in SST
at the buoy was also associated with the rapid southward displacement of the
NBF.
The tramontane and mistral are two strong, dry and cold northwesterly/northerly winds which are channelled by the topography, between
the Pyrenees and the Massif Central for the tramontane and between the
Massif Central and the Alps for the mistral (Fig. ). These winds
produce strong air–sea exchanges which lead to
important diabatic buoyancy losses for the sea and which also impact the NBF
position in autumn . Furthermore, the tramontane and
mistral have up- or down-front wind components over the NBF and the Northern
Current. Down- (up-)front wind refers to a wind parallel and in the same
(opposite) direction as the current (Fig. ). In an academic study,
showed that the up- or down-front wind can, respectively, re-stratify or de-stratify the mixed layer. This modification of stratification
is due to the differential advection of density by the Ekman transport, which
destabilises the front and releases the symmetric instability
. To represent the modification of stratification generated
by the Ekman transport, defined a surface adiabatic
buoyancy flux named the Ekman buoyancy flux as the product between the Ekman
transport and the horizontal surface buoyancy gradient. This physical
diagnostic was applied to the real case of the Kuroshio and
Gulf streams , showing the major role of the Ekman buoyancy
flux as compared to the diabatic buoyancy flux in the modification of
stratification in frontal zones. In the NWMS, a realistic modelling study
also analysed the respective roles of the diabatic
air–sea and adiabatic Ekman buoyancy losses along the Northern Current front
on the deep-water formation during winter. This study found that the Ekman
buoyancy flux contribution was dominant along the northern branch of the
Northern Current.
In the NWMS , , , and showed
the importance of using air–sea coupled modelling to represent the rapid
variations in air–sea exchanges induced by the decrease in SST during
tramontane and mistral events. Furthermore, in frontal zones, the air–sea
exchanges are clearly an air–sea coupled process. The SST front and sea
surface frontal currents directly impact the air–sea exchanges with
significant feedback between the sea and the atmosphere .
The atmospheric response to the front creates wind stress divergence and curl
impacting the sea by Ekman pumping . These air–sea
feedbacks in frontal zones may also have a marked influence on the diabatic
and Ekman buoyancy fluxes .
The first aim of this paper is to describe the dynamics of the NBF during
IOP16 of the HyMeX programme. Our approach combines satellite data, in situ
data and air–sea coupled simulation at a kilometre scale. A second aim is to
evaluate the air–sea exchanges in the frontal zone and investigate their
impact on ocean stratification. In particular, the competing roles of the
diabatic buoyancy flux, the adiabatic Ekman buoyancy flux and advective
processes will be assessed by means of an original stratification budget
diagnosis.
This paper is organised as follows. Section 2 presents the IOP16 case study.
In Sect. 3, the results of the air–sea coupled simulation are analysed and
discussed with respect to the available observations. Then, the NBF dynamics
and associated air–sea exchanges are investigated in Sects. 4 and 5,
respectively. The computation and results of the stratification budget
diagnosis are presented in Sect. 6. Finally, the results are discussed and
some conclusions are drawn in Sect. 7.
Case studyAtmospheric conditions
Figure shows the mean sea level pressure (MSLP) and the surface
air temperature and wind from the European Centre for Medium-range Weather
Forecasts (ECMWF) analysis at 12:00 UTC for IOP16a (26 October 2012) and
IOP16b (27, 28 and 29 October 2012). IOP16a (Fig. a) was
characterized by low pressure centred over the Pyrenees and a southwesterly
wind advecting warm air (temperature above 20 ∘C) over the NWMS. On
the next day (Fig. b), the low pressure was positioned over the
Alps, leading to flow reversal with northwesterly/northerly winds advecting
cold air over the NWMS. On 28 October (Fig. c) the surface low was
deeper and located over the Ligurian Sea. The northwesterly/northerly wind
increased and extended over the western Mediterranean basin with an air
temperature lower than 10 ∘C in the NWMS. Finally, on
29 October (Fig. d) the cyclone dissipated and the air temperature
increased again.
Atmospheric surface synoptic
conditions from ECMWF analyses at 12:00 UTC (a) on 26 October,
(b) on 27 October, (c) on 28 October and (d) on
29 October 2012. Coloured area: atmospheric surface temperature in
∘C. Contour lines: mean sea level pressure in hPa. Arrows:
surface wind. The black triangle indicates the position of the Lion
meteorological buoy.
The Lion meteorological buoy, positioned in the abyssal plain of the Gulf of
Lion (42.06∘ N–4.64∘ E, indicated by the black triangle in
Fig. ) and in the tramontane and mistral flows, provides hourly
measurements of the 10 m wind, the 2 m temperature and humidity, and the
SST (Fig. ). During IOP16a (26 October 2012) the Lion buoy measured
low wind speeds between 5 and 10 ms-1. The 2 m temperature and
relative humidity remained nearly constant around 20 ∘C and
80 %, respectively. At 09:00 UTC a brief decrease in wind and 2 m
temperature was observed, corresponding the signature of a precipitation
event. Then, on 27 October, the wind speed increased from 8 ms-1
at 06:00 UTC to a maximum of 26.2 ms-1 at 18:00 UTC. On
28 October, the wind speed remained nearly constant around
22 ms-1 and finally decreased on 29 October to less than
5 ms-1 at the end of the day. A rapid decrease in the 2 m
temperature was observed on 27 October during the wind increase, with a
temperature drop of 7.5 ∘C in a few hours. After this rapid
decrease, the 2 m temperature continued to fall slowly, reaching a minimum
of 8 ∘C on 28 October at 09:00 UTC before rising slowly by
3 ∘C on 29 October. During IOP16b, the 2 m humidity decreased from
80 % at 00:00 UTC on 27 October to 40 % at 00:00 UTC on
30 October.
Time series at the Lion meteorological buoy of (a) the
10 m wind speed, (b) the 2 m air temperature, (c) the
2 m air humidity and (d) the sea surface temperature. Observations
in black and simulations in blue.
To summarise, this autumn tramontane and mistral storm (27, 28 and
29 October 2012) was characterised by northerly/northwesterly winds exceeding
20 ms-1. Meanwhile, the surface air temperature and humidity
dropped by 10 ∘C and 40 %, respectively.
Oceanic conditions
Figure a–f shows two SST analysis products from satellite
observations before and after the strong-wind event, and the SST
difference.The first analysis is the global OSTIA product ,
with a horizontal resolution of about 6 km. This product is used in the
ECMWF operational model. The second analysis is the Mediterranean Copernicus
product provided at higher resolution (about 1 km).
The spatial resolutions indicated above refers to the resolution at which the
data are provided but not to their effective resolutions. There is no
significant difference between these two data sets. The NBF extends over
several tens of kilometres in both products with steeper gradients in the
high-resolution product. At 00:00 UTC on 25 October (Fig. a and d)
the NBF separated the northern water at 18.5 ∘C from the southern
water at 21.5 ∘C, yielding a surface thermal front of 3 ∘C.
The Lion buoy was positioned in the warm side of the frontal zone. At the end
of IOP16 (30 October 2012, Fig. b and e) the NBF had moved several
tens of kilometres to the south and the Lion buoy was then positioned in the
cold side of the frontal zone. During IOP16, the SST decrease
(Fig. c and f) was about 2 ∘C in the basin and reached
more than 3 ∘C around the Lion buoy. The temperature of the southern
water (northern water) was around 16 ∘C (20 ∘C),
establishing a surface thermal front of 4 ∘C. At the Lion buoy, the
SST in situ measurement (Fig. d) showed a rapid decrease in SST of
about 4 ∘C between 12:00 UTC on 27 October and 00:00 UTC on
29 October in agreement with the analyses.
Sea surface temperature analysis from the OSTIA product
(a–c), from the Mediterranean Copernicus product (d–f), from the coupled simulation (g–i) at 00:00 UTC on
25 October (a, d and g) and at 00:00 UTC on
30 October (b, e and h), and the difference between
00:00 UTC on 30 and on 25 October (c, f and i).
The triangle indicates the Lion buoy. The dots correspond to the position of
the Argo profiles shown in Fig. and the dashed lines to the glider
sections shown in Fig. .
Gliders provided information on the oceanic stratification and the spatial
distribution of water masses during IOP16 (see their position in
Fig. ). Two gliders were present near the Lion buoy on
25 October 2012, in the frontal zone (Fig. ). During the IOP, the
first glider, named Eudoxus, moved to the west (Section A–B in
Fig. ). The second glider, named Campe, moved to the south
(section C–D in Fig. ). Before the strong wind, according to the
two gliders, the mixed layer depth (MLD) near the Lion buoy was about
40 m, the potential temperature was higher than 19.5 ∘C in
the mixed layer (Fig. d and f) and the salinity was lower than 38.2
(not shown). This temperature corresponds to the surface temperature of the
cold side of the NBF (Fig. a and d). According to the Eudoxus
glider (Fig. d), between the two ends of the section, the
difference in the MLD was about 60 m and the temperature drop about
3.5 ∘C. These differences resulted from both the crossing of the NBF
and the onset of the strong wind. According to the Campe glider
(Fig. f), which remained on the southern side of the front, the
difference in the MLD was 40 m and the temperature increased by
0.5 ∘C. However, along this glider track, the 2 ∘C
temperature drop measured between 27 and 29 October was probably due to the
strong-wind event whereas the temperature increase observed after
29 October resulted from the southward displacement of the glider (see
Fig. e).
SST measurement (a), simulation (b) and bias of
the simulation (c), along the trajectory of the glider Eudoxus
(moving westward) and Campe (moving southward). Vertical section of potential
temperature between 0 and 200 m depth for the Eudoxus glider
observation (d) and simulation (e) and the Campe glider
observation (f) and simulation (g).
Furthermore, two Argo profiles (see their position in Fig. ) were
available during IOP16 in the frontal zone (Fig. ). The first
profile was obtained on 26 October, 40 km east of the Lion buoy
(41.8∘ N–5.05∘ E), and the second one on 31 October,
30 km east of the Lion buoy (41.8∘ N–4.86∘ E). In
the mixed layer, the potential temperature drop between the two profiles was
large (about 5 ∘C) while the salinity decrease was about 0.05.
Finally, the potential density increase (about 1.32 kgm-3 at the
surface) was driven directly by the potential temperature decrease and the
MLD deepened from 25 to 40 m.
Argo potential temperature (a), salinity (b) and
potential density (c) measured (solid line) and simulated (dashed
line) at the different dates indicated in the figure. For 31 October, only
simulated profiles are given (in green). The horizontal lines (c)
indicated the MLD for each density profiles.
In summary, the observations showed a rapid decrease in the SST (greater than
4 ∘C) in the frontal zone. This rapid decrease was associated with
the southward displacement of the NBF and a deepening of 20 m of the MLD in
the frontal zone. As satellite and in situ data only give a partial view of
the NBF dynamics, and in order to better describe and understand this
process, a high-resolution air–sea coupled simulation was performed.
Numerical experimentAir–sea coupled simulation
To perform the simulation, we used the Meso-NH–SURFEX–Symphonie system
(Meso-NH: , SURFEX: , Symphonie:
) based on the SURFEX–OASIS
interface . This coupled system was described and
validated in .
A high-resolution configuration of this coupled system was implemented with a
horizontal resolution of 1 km in the oceanic model and of
2.5 km in the atmospheric model. At this resolution, the oceanic
model can be considered as eddy-resolving. In the NWMS, the Rossby radius is
of the order of 5 to 10 km , which corresponds to the
effective resolution of the oceanic model. showed the
ability of the Symphonie oceanic model to reproduce the symmetric instability
process at this resolution. However, at this resolution, the oceanic model
convection must be parameterised, as described in . The
non-solar heat fluxes and water fluxes are actually redistributed over the
whole mixed layer, in accordance with . The atmospheric
model at 2.5 km can be considered as a convection-permitting model; i.e. the convection is directly resolved by the atmospheric model equations.
In the vertical, the oceanic model uses 40 generalised sigma vertical levels,
10 of them in the first 100 m (above the abyssal plain) with a
resolution just below the sea surface of 1.5 m. The atmospheric model uses
52 terrain-following vertical levels ranging from 15 to 15 000 m.
Using the SURFEX OASIS3-MCT interface , the SST and sea
surface current are sent to the SURFEX model, which then returns the wind
stress (τ), the shortwave flux (SW), the non-solar heat fluxes (Qns=LW+H+L with LW being the longwave flux, H the sensible heat flux
and L the latent heat flux) and the water fluxes (E-P, evaporation minus
precipitation). The turbulent fluxes are calculated using the bulk
parameterisation developed by . The coupling frequency is set
to 10 min. Flux calculations are carried out on the atmospheric model grid
(after bilinear interpolation of the oceanic fields) and are then bilinearly
interpolated on the grid of the ocean model.
The atmospheric model covers the whole western Mediterranean basin (same
spatial coverage as in ). The ocean model covers the
western Mediterranean Sea, excluding the Alboran Sea and part of the
Tyrrhenian Sea (same grid as in ). In the oceanic part
of the atmospheric grid not covered by the ocean grid, the air–sea fluxes are
computed using the sea surface temperature provided by the SST OSTIA product
and without using sea surface current.
The coupled simulation started at 00:00 UTC on 25 October 2012 (24 h before
the beginning of IOP16) and ended at 00:00 UTC on 30 October 2012 (at the
end of IOP16). The ocean model was initialised with a coupled simulation as
described in (from all the simulations presented in this
paper, the so-called MOON simulation was chosen for its better agreement
with the available observations). Spin-up was avoided as the initial state
given by the MOON simulation was based on the same oceanic model
configuration and grid. The boundary conditions of the ocean model were
provided by the analyses of the MERCATOR-OCEAN operational centre
based on the NEMO model . Initial
and boundary conditions of the atmospheric model were provided by the 6 h
ECMWF analyses with a horizontal resolution of 1/8∘.
Simulation validation
The simulation was validated with in situ and satellite observations. At the
Lion buoy (Fig. a, b and c), the simulation correctly reproduced
the wind intensity evolution but the wind absolute maximum at 18:00 UTC on
27 October was underestimated by 5 ms-1. The 2 m temperature
and humidity were also well represented, but the rapid decrease in
temperature observed around the wind maximum was slower in the model and a
positive humidity bias of 10 % appeared at the end of the
simulation.
Then, to validate the SST decrease in the NWMS and the NBF dynamics, the SST
satellite analyses and the simulated SST were compared (Fig. ). The
simulated temperature field contained more small-scale structures (fronts,
eddies, filaments) than the corresponding satellite field. The horizontal
resolution of the latter was, in fact, substantially lower than that of the
model. The initial surface temperature and NBF position were well represented
by the simulation, but the extension of the warm AW to the east of the Lion
buoy was greater than in the observations. At the end of IOP16, the SST
decrease and the NBF southward shifting were also well represented even if
the latter was less pronounced than in the observations. At the Lion buoy
(Fig. d), the SST decrease of 4 ∘C was well represented by
the simulation but the decrease started 6 h earlier.
Finally, to validate the evolution of oceanic stratification, the simulation
results were compared to the profiles obtained with the gliders and Argo
(Figs. and ). The comparison with the gliders showed
that the initial state of the model was satisfactory (Fig. ) near
the Lion buoy even though the MLD was less deep (by about 20 m) and
the thermocline thicker, probably because of the insufficient vertical
resolution of the model. In comparison with results from the first glider
(Eudoxus), which moved westward, the SST and MLD deepening were well
represented. However, at the end of the section, the MLD was deeper than in
the observation, probably because of the crossing of the Northern Current
visible from the slope of the isopycnals, which differed in the simulation
and the observations. For the second glider (Campe), which moved southward,
the temperature evolution was represented differently by the glider and the
simulation due to the presence of small-scale structures in the simulation
that were not observed. During strong-wind events, the temperature decrease
was less intense in the simulation but the MLD deepening was well
represented. Finally, after the strong-wind event (29 October), the
temperature increase was not simulated. The Argo comparison (Fig. )
showed good agreement for the initial temperature profile, while a surface
negative salinity bias of about 0.2 was shown by the model. At the end of the
strong-wind event, the simulation showed a decrease in temperature and an
increase in salinity, density and MLD. The temperature decrease was smaller
than in the observations – 3.3 ∘C compared to 5.3 ∘C – whereas the salinity increase was more important in the simulation than in
the observations: 0.15 compared to 0.05. The increase in density and MLD were
smaller than in the observations: 1 kgm-3 and 15 m
compared to 1.32 kgm-3 and 25 m, respectively. However,
it is worth noting that the observed profile is located to the north of the
NBF, whereas the corresponding one in the simulation is within the frontal
zone. When the comparison is made with a simulated profile located slightly
further east (41.8∘ N–5.15∘ E, green dashed line in
Fig. ), the simulation results are much closer to the Argo
observations.
To conclude, this simulation is satisfactory regarding the atmospheric and
oceanic evolution, including air and sea parameters at the surface, the
oceanic mixed layer deepening, and the NBF southward shift during the strong-wind event. In addition, the 1 km resolution oceanic simulation
correctly represents the narrow surface front and associated meso- and
sub-mesoscale structures, the position of which differs slightly from the
actual ones.
Characteristics and dynamics of NBF
The NBF dynamics during IOP16 (26 to 30 October 2012) are now examined through
an analysis of the air–sea coupled simulation.
Figure a–b represents the surface potential density spatial
distribution before (00:00 UTC on 26 October), and after (00:00 UTC on
30 October) the IOP16 and Fig. c shows the density difference
between these two dates. Two specific surface water masses of the NWMS (the
relatively light AW to the south and the dense AW to the north) appear clearly in Fig. a and b (in blue and yellow–red, respectively).
Before IOP16 (Fig. a), the Atlantic waters protrude northward
forming a meander of light water extending to 42∘ N and from 3 to
6∘ E. The difference in surface density between the two sides of the
front is about 1 kgm-3 with a surface density gradient locally
greater than 0.1 kgm-3 per kilometre. Filaments formed at the
periphery of the meander extend northward. Further to the east (around
6.5∘ E) a second intrusion of light water occurs, but its northward
extension is smaller and its density front weaker. These two meanders shape
the NBF, which can be characterised by the 27.0 kgm-3 surface
isopycnal. After IOP16 (Fig. b), the light water patch is shifted
to the southwest and the surface potential density increases in the NWMS. The
NBF surface density gradient is reduced except along the western part of the
meander, and filament structures have dissipated. The NBF is now
characterised by the 28.0 kgm-3 surface isopycnal. Looking at
the surface density evolution between the beginning and the end of IOP16, the
maximum potential density increase (Fig. c), larger than
1.5 kgm-3, is located in the area impacted by the NBF dynamics
(along the meander) and along the north coast of the Gulf of Lion. Note the
appearance of cold, dense waters along the French coast (around
5.5∘ E), a sign of coastal upwelling.
Sea surface density in kgm-3(a, c, e) and
stratification index at 250 m in kgm-2(b, d, f) at
00:00 UTC on 26 October (a, b) and at 00:00 UTC on
30 October (c, d) and the difference between 00:00 UTC on 30 and
26 October (e, f). The black contour (a, e) represents the
27.0 kgm-3 isopycnal, the blue contour (c, e)
represents the 28.0 kgm-3 isopycnal and the red line (a, c) represents the surface density gradient greater than
0.1 kgm-3 per kilometre. The straight black lines indicate the
position of the vertical sections shown in Fig. .
The cumulative effect of the strong-wind event on the stratification of the
upper layers can be analysed through the stratification index (SI)
.
SI(H)=∫H0(ρ(H)-ρ(h))dh,
where ρ is the potential density (kgm-3) and H the
reference level (m). SI is expressed in
kgm-2. It represents the amount of buoyancy to be extracted to
mix the water column from the surface to level H and achieve a homogeneous
density ρ(H).
Figure d–f shows the SI at 250 m (a depth below which no
significant changes are observed) before and after IOP16 and the SI
differences between these two periods. Before the IOP16 (Fig. d),
the light Atlantic water corresponds to higher stratification, with an SI at
250 m greater than 120 kgm-2. Inside this meander, a mesoscale
anticyclonic eddy has a noticeable effect on stratification with an SI higher
than 200 kgm-2 (around 41∘ N and 4.5∘ E). The
NBF separates this patch of highly stratified water from the less stratified
Atlantic water characterised by an SI lower than 90 kgm-2, i.e.
a stratification front of about 30 kgm-2. The filaments visible
in the surface potential density map (Fig. a) are too thin to have
a significant impact on the upper layer integrated stratification. The SI
evolution during IOP16 (Fig. f) shows a maximum loss of
stratification (about 50 kgm-2) in the northern and eastern
parts of the NBF meander and a stratification gain (about
20 kgm-2) in the western part of the NBF meander. Outside the
frontal region, SI variations appear to be related to the displacement of
mesoscale and sub-mesoscale eddies.
Two vertical sections of potential density are now examined to illustrate the
vertical processes. Figure presents the vertical distribution of
isopycnals before (black lines in Fig. ) and after IOP16 (blue
lines in Fig. ) and the seawater potential density difference (in
colours) along north–south (NS) and east–west (EW) sections (indicated by
black lines in Fig. ).
(a) South–north section at 4.65∘ E and
(b) west–east section at 41.7∘ N (see Fig. for
their positions) of the potential density difference between 00:00 UTC on
30 October and 00:00 UTC on 25 October. Contour lines: Isopycnals on 25
(black lines) and 30 October 2012 (blue lines).
On the NS section (Fig. a), before the strong-wind event, the
section intersects the NBF meander at 42.1∘ N (the NBF position is
defined by the 27.0 kgm-3 isopycnal). South of the NBF, a patch
of light AW, with a density lower than 27.0 kgm-3, extends over
a depth of about 50 m. To the north, the light water thins to
10 m around 42.4∘ N, corresponding to the filament
structure. This section also intersects the mesoscale anticyclonic eddy
previously described around 41∘ N (explaining the local deepening of
the isopycnals) and the Northern Current along the bathymetric slope.
Between 41.75 and 42.75∘ N, the doming of the isopycnals corresponds
to the cyclonic gyre. After the strong-wind event, the NBF migrates southward
by about 0.3∘ (about 35 km), and the NBF position is now
defined by the 28.0 kgm-3 isopycnal. The maximum increase in
potential density (more than 1.5 kgm-3 to a depth of
40 m) is obtained in the displacement area of the NBF.
The EW section (Fig. b) crosses the meander (Fig. c)
before IOP16. It intersects the NBF twice, at about 4.0 and 6.0∘ E.
The isopycnal doming, corresponding to the cyclonic gyre, is also present
between 3.4 and 4.4∘ E, while the dive of the isopycnals to the west
indicates the change in direction of the Northern Current, which follows the
isobaths of the continental slope. After IOP16, the eastern intercept point
of the NBF with the section of Fig. b has been shifted westward by
about 1.0∘ (about 80 km). The western intercept point has
also moved westward but only by about 0.2∘ (about 16 km). The
maximum increase in potential density is obtained in the eastern part of the
NBF, with an increase greater than 1.0 kgm-3 to a depth of
30 m. The density increase in the core of the Atlantic water (between
3.8 and 5.2∘ E) is smaller (<1kgm-3 at
40 m). Near the western part of the NBF (around 3.7∘ E), the
density increases at the surface and decreases at 50 m depth due to
the isopycnal deepening, which explains the SI increase in this region
(Fig. f).
To conclude, during the strong-wind event, a horizontal displacement of the
NBF of several tens of kilometres is simulated. This directly impacts the
evolution of oceanic stratification in the NBF zone. The stratification
evolution is also impacted by vertical advective processes producing
isopycnal deepening.
Air–sea exchanges
The air–sea exchanges can be divided into three fluxes: the wind stress
(τ, in
Nm-2), the net heat flux
(Qnet=SW+LW+H+LE, in Wm-2)
and the water flux (Fw=E-P, in mmh-1).
Figure shows the time evolution of wind stress, net heat flux and
water flux calculated at the Lion buoy, positioned in the NBF zone. The wind
stress variation (Fig. a) corresponds to the wind speed variation
(Fig. a), with a maximum wind stress higher than
1.5 Nm-2. On average, the net heat flux (Fig. b) was
close to zero between 25 and 26 October. During this low-wind period the
solar heat flux balanced the non-solar heat flux. During the strong-wind event, upward (i.e. cooling the sea) net heat fluxes were correlated with the
wind speed, with a maximum net heat flux close to 2000 Wm-2 on
28 October 2012. The water flux (Fig. c) was impacted by the high-precipitation event occurring on 26 October, with a maximum downward
instantaneous water flux of about 25 mmh-1. During the strong-wind event, evaporation dominated and led to an upward (positive) water flux.
The spatial distribution of air–sea fluxes is now examined.
Figure represents the wind stress, net heat flux and water flux
averaged over IOP16a (26 October 2012) and IOP16b (27–29 October 2012).
During IOP16a the NWMS is dominated by weak southwesterly winds
(Fig. a). During this period, the NBF directly impacts the wind
stress distribution. The maximum wind stress appears to the south of the
front over the patch of warm AW. The net heat fluxes (Fig. c) are
also directly impacted by the NBF position. The maximum heat loss appears on
the warm patch, due to the higher SST, directly impacting turbulent heat
fluxes. Downward water fluxes (Fig. e) are organised in bands
during this high-precipitation event. During IOP16b, the NWMS is dominated by
strong northerly/northwesterly winds (Fig. b), corresponding to
tramontane and mistral events. The maximum wind stress is located in these
wind veins, off the Gulf of Lion. The NBF has no clear impact on the wind
stress spatial distribution. On the other hand, the net heat flux and water
flux are directly impacted by the NBF position, with a maximum heat and water
loss higher than 1200 Wm-2 and 30 mmday-1 on the
warm side of the front, and cross-frontal differences greater than
200 Wm-2 and 10 mmday-1.
Time series of (a) the wind stress, (b) the net
heat flux and (c) the water flux simulated at the Lion
meteorological buoy.
The wind stress curl during IOP16a and IOP16b is presented in Fig. . As explained by , surface winds are
weaker on the cold side of a thermal front. This produces a divergence when
the wind blows across the front and a curl when the wind blows parallel to
the front. In the case of the NBF meander during the low-wind period of
IOP16a, Fig. a indicates the wind curl modifications linked to the
front. In the eastern part of the meander, a negative along-front wind stress
curl appears clearly, and, in the western part of the meander, a positive
along-front wind stress curl also appears clearly. During IOP16b the wind
stress curl appears principally around the tramontane and mistral wind
corridors. A positive (negative) wind stress curl corresponds to the left
cyclonic (right anticyclonic) side of the wind vein. In the eastern part of
the NBF meander (around 5∘ E and 41.8∘ N) a positive wind
stress curl appears. However, it is difficult to know whether the wind stress
curl is directly connected to the NBF meander or to the fine wind jet
structures. Finally, during the strong-wind event, unlike the situation in
the low-wind event, the NBF meander does not appear to have a major influence
on the wind stress curl.
(a, b) Wind stress (Nm-2); (c, d) net
surface heat flux (Wm-2); (e, f) surface water flux
(mmday-1) averaged over IOP16a (a, c, e) and IOP16b
(b, d, f). Contour lines: sea surface density averaged over IOP16a
(a, c, e) and IOP16b b, d, f; contour interval:
0.3 kgm-3).
Wind stress curl (Nm-3×106) averaged over
IOP16a (a) and IOP16b (b). Contour lines: sea surface
density averaged over IOP16a (a) and IOP16b (b); contour interval:
0.3 kgm-3.
To conclude, the NBF position and dynamics directly impact the spatial and
temporal distributions of air–sea heat and water exchanges. Furthermore,
during the low-wind period, the NBF also impacts the wind stress with the
generation of along-front wind stress curl.
Stratification budget diagnosis
The stratification variation in water column can be modified through diabatic
processes and horizontal or vertical advection of buoyancy. Most models
assume that the stratification budget is essentially driven by
one-dimensional turbulent mixing of heat and water at the air–sea interface.
Following , the air–sea exchanges induce a surface or
diabatic buoyancy flux that can be diagnosed as
B0=gαQnetρ0Cp+gβSSS(E-P),
where B0 is the diabatic buoyancy flux in m2s-3, α the
thermal expansion coefficient in K-1, Cp the specific
heat capacity in Jkg-1K-1, β the saline contraction
coefficient, SSS the sea surface salinity, E the evaporation and P the
precipitation in ms-1.
This one-dimensional approximation is suitable if the ocean is horizontally
homogeneous. In reality, mesoscale and sub-mesoscale structures populate the
ocean. Theses structures are marked by horizontal buoyancy fronts. As shown
by and , the stratification can be
significantly modified by interactions between these fronts and Ekman flow
generated by frictional forcing. When the winds are down-front, the density
advection of dense water over light water by Ekman transport destabilises the
water column and triggers convection. This process de-stratifies the water
column by Ekman advection of buoyancy and mixing through the mixed layer. On
the contrary when the winds are up-front the Ekman flow yields an Ekman
advective re-stratification in the surface layer. Following
, the frictional forcing induces a wind-driven or Ekman
buoyancy flux (EBF), which can be diagnosed as
EBF=-gρ0Me⋅∇hρ(z=0),
where the EBF is in m2s-3, Me is the Ekman transport (Eq. )
in m2s-1, ∇h is the horizontal gradient, g is the
gravitational acceleration and ρ(z=0) is the surface density in
kgm-3.
Me=z^×τρ0ζa,
where z is the vertical unit vector, τ the wind stress in Nm-2, and
ζa the absolute vorticity in s-1.
The stratification index variation at depth H (with H=250 m, H> MLD)
between times T1 and T2 can be approximated by the integral of buoyancy mass
flux between times T1 and T2 (Eq. 5). In order to evaluate the competing
roles of the diabatic and Ekman buoyancy fluxes on stratification variation,
these two term are diagnosed and compared to the stratification variations.
Finally, to close our stratification budget diagnosis we evaluate the
residual term corresponding to other potential sources of horizontal and
vertical advection of buoyancy (geostrophic circulation, frontogenesis, Ekman
pumping, etc.) that are not directly diagnosed in this study.
ΔSI(H)=SI(H)T2-SI(H)T1=ρ0g∫T1T2B0(t)⋅dt+ρ0g∫T1T2EBF(t)⋅dt+R,
where the first and second right-hand-side terms represent the mass fluxes, induced by diabatic and friction
processes, respectively, whereas the third term induces all the remaining
processes.
Figures and present the spatial distribution of the
different terms of the stratification budget diagnosis (Eq. ) for
IOP16a and IOP16b. During the low-wind period (IOP16a), the stratification
(Fig. a) decreases (increases) by about 20 kgm-2 along the western (eastern) part of
the NBF meander. This stratification
evolution is not directly controlled by the diabatic and Ekman buoyancy mass
fluxes (Fig. b and c), which are small relative to the residual
term (Fig. d). The advective processes play a dominant role in the
evolution of stratification during this period of low wind. During the strong-wind event (IOP16b), in contrast to the low-wind period, the stratification
(Fig. a) decreases (increases) by about 60 kgm-2 (20 kgm-2) along the eastern (western) part of
the NBF meander. The
evolution of the stratification (Fig. a) is directly controlled by
the diabatic and Ekman buoyancy mass fluxes (Fig. b and c). The
cumulated diabatic buoyancy mass flux loss (Fig. b) is between 15
and 30 kgm-2 in the tramontane and mistral corridor, being
slightly larger south of the NBF (not shown in Fig. b but visible
for the net heat flux in Fig. c). However, stratification
evolution in the NBF meander zone is principally driven by the Ekman buoyancy
mass flux (Fig. c). In the east (west) of the NBF meander, a
down-front (up-front) wind generates an Ekman buoyancy flux of -30 to
-45 kgm-2 (15 to 30 kgm-2). In a second order,
the residual term (Fig. d) also plays a role in how stratification
evolves in the NBF zone, particularly in the western part of the NBF meander,
with a stratification gain.
Stratification variation budget diagnosis (Eq. ) during
IOP16a: (a) SI difference, (b) diabatic buoyancy mass flux,
(c) Ekman buoyancy mass flux and (d) residual term. Red
circles indicate the areas chosen for the budget diagnostic in
Fig. .
Same figure as Fig. for IOP16b.
Figure shows the time evolution of the different terms of the
stratification budget (Eq. ) in the eastern and western parts of
the NBF meander. Firstly, the presence of oscillations produced by inertial
waves can be noted on the evolution of the stratification. These oscillations
have a period of about 1 h and an amplitude of about 10 kgm-2.
Except for these oscillations, the stratification remains approximately
constant from 25 to 27 October (IOP16a). During IOP16b, the stratification
decreases (increases) from 122 to 69 kgm-2 (from 86 to 102 kgm-2) along the eastern (western) part of the NBF meander. Along
the east of the NBF meander (Fig. a), the stratification decrease
(53 kgm-2) is due to the buoyancy fluxes and, more precisely, the Ekman buoyancy flux, (45 kgm-2 compared to
17 kgm-2 for diabatic buoyancy mass flux). In this region, the
advective processes remain negligible (about 1 kgm-2). Along the
west of the NBF meander (Fig. b), the stratification decreases by
about 5 kgm-2 during the low-wind period and increases by about
20 kgm-2 during the strong-wind period. The stratification
decrease during the low-wind period is driven by a negative Ekman buoyancy
flux. During the strong-wind period, the Ekman buoyancy flux of
20 kgm-2 balances the diabatic buoyancy loss of
18 kgm-2. The residual term of about 18 kgm-2 plays
a major role in the stratification increase in this zone.
Time series of the different terms of the stratification budget
diagnosis (SI in black, diabatic buoyancy flux in blue, Ekman buoyancy flux
in red, total buoyancy flux in magenta and residual in green): (a)
for the eastern part of the NBF meander (averaged over 10 km2 around
41.9∘ N and 5.1∘ E) and (b) for the western part
of the NBF meander (averaged over 10 km2 around 41.6∘ N and
3.9∘ E; see Figs. and ). In order to compare
the different terms, all the curves start with the same value, which is the
initial SI.
The residual term is not negligible in our stratification budget diagnosis,
particularly along the western part of the NBF meander. A source of this term
is the vertical advection generated by the Ekman pumping directly connected
to the wind stress curl (Eq. ).
We=∇×τρ0ζa
To illustrate this, Fig. shows the 28.5 kgm-3
isopycnal depth variation and the Ekman pumping averaged over IOP16a and
IOP16b. During IOP16a, the isopycnal depth variation (Fig. a) is
directly anti-correlated with the residual term of the stratification budget
(Fig. d). The stratification decreases in ascending isopycnal
zones and increases in subsiding isopycnal zones. This anti-correlation
illustrates the major role of the vertical processes in the residual term.
The Ekman pumping, generated by the interaction between the wind and the
front, is an important source of this vertical advection (Fig. c).
Although, at the scale of the whole domain, the Ekman pumping does not appear
to be correlated with the change in the level of the isopycnals, this is not
the case along some parts of the meander such as its western part, where the
angle between the wind and the isopycnals is small and the Ekman pumping
seems to contribute to the change in stratification. During IOP16b, the
isopycnal depth variations (Fig. b) are also anti-correlated with
the residual term of the stratification budget (Fig. d). In the
western, unlike in the eastern, part of the NBF meander, the isopycnal
deepening is clearly produced by the Ekman pumping (Fig. d).
Another source of vertical advection in frontal zones is the ageostrophic
secondary circulation cell generated by down-front wind and frontogenesis
processes , with upward and downward vertical
currents, respectively, in the light and dense parts of front. It is probably
this circulation that impacts the isopycnal depth in this zone.
Panels (a, b) show 28.5 kgm-3 isopycnal deepening (m), and
(c, d) Ekman pumping (m) during IOP16a (a, c) and IOP16b
(b, d). Contour lines: sea surface density averaged over IOP16a
(a, c) and IOP16b (b, d). Contour
interval: kgm-3.
In conclusion, to a first-order approximation, the stratification variations
in the eastern part of the NBF (down-front wind) during IOP16b are
principally driven by the Ekman buoyancy flux generated by the interactions
between the strong-wind event and the NBF. However, the residual term of
stratification budget diagnosis is not negligible, particularly in the
western part of the front through an Ekman pumping effect.
Summary and discussion
This case study focused on the evolution of NBF dynamics and stratification
during IOP16 (26–29 October 2012), which, during a first period (IOP16a),
was characterised by weak southwesterly winds followed, in period IOP16b, by
a strong northerly wind event.
Before IOP16, the light AW present south of the NBF spread up to
42∘ N. After IOP16, the observations showed a southward displacement
of the NBF, of several tens of kilometres, with a rapid decrease in SST,
larger than 4 ∘C in the frontal zone. To better describe and
understand the NBF dynamics during this case study, we performed an air–sea
coupled simulation at a kilometre scale. This simulation is in good agreement
with the observations and has the ability to reproduce meso- and sub-mesoscale
structures (eddies, fronts, filaments). The simulation showed that, before
the strong northerly wind event, the NBF reached the Lion buoy and surface
filaments of light water (not visible in the satellite SST) became detached
and were entrained northward. During the strong-wind event, the simulation
reproduced the rapid displacement of NBF, of several tens of kilometres, and
the SST decrease of 4 ∘C at the Lion buoy. Furthermore, the
filaments were dissipated. As suggested by , the SST
decrease at the Lion buoy was associated with the NBF dynamics. These dynamics
also impacted the oceanic stratification with a significant loss (gain) of
about 40 kgm-2 (20 kgm-2) of stratification along
the eastern (western) part of the NBF meander.
During the period of light southwesterly wind, the NBF clearly impacted the
wind stress and the net heat flux through two effects: first a decrease in
the air–sea exchanges on the cold side of the NBF (to the north) and second,
as explained by , the formation of a wind stress curl.
During the strong-wind event, contrary to the low-wind event, the NBF had
little impact on the wind stress distribution. The wind stress curl was
mainly linked to the channelling of the tramontane and mistral by the
continental orography. In contrast, the NBF directly impacted the heat and
water budgets. The rapid SST cooling induced significant air–sea coupling, as
suggested in other case studies in the NWMS
, by
stabilising the unstable atmospheric boundary layer and leading to a decrease in the turbulent fluxes. Furthermore, this study pointed out that the surface
cooling was partly associated with a rapid NBF displacement, which led to
an additional air–sea coupling feedback. Therefore, a good representation of the
frontal dynamics is essential to correctly depict the space and time
evolution of the air–sea exchanges.
Finally, a stratification budget diagnosis was performed during the low- and strong-wind periods. The results differed greatly according to the wind
regime. During the low-wind period, the evolution of the stratification in
the frontal zone was directly controlled by the advective processes whereas,
during the strong-wind period, it was controlled by the Ekman buoyancy flux,
which could be up to 3 times stronger than the diabatic heat flux in the
eastern part of the NBF. This flux directly impacted the frontal dynamics and
the stratification variations. The stratification increased by up to
20 kgm-2 in the western part of the NBF meander (where the wind
was parallel to and in the same direction as the current) and decreased by up
to 40 kgm-2 in the eastern part of the NBF meander (where the
wind was parallel but in the opposite direction to the current). The
advective processes also played a role in the evolution of stratification.
One of these processes clearly appears to be the Ekman pumping. Finally, the
inertial waves also impacted the stratification, with variations of the order
of 10 kgm-2.
To the best of our knowledge, this study is the first realistic study that
uses an air–sea coupled model to evaluate the impact of Ekman buoyancy fluxes
on an oceanic front that is not topographically controlled. Without
bathymetric constraint, the front moves by several tens of kilometres during
the strong-wind events. This displacement is associated with a marked loss of
buoyancy and a rapid de-stratification, largely induced by the wind–front
interaction. In an academic study, described a
frontogenesis process generated by the Ekman buoyancy flux. This process is
not reproduced in our study probably because of the absence of bathymetric
constraint and front displacement. On the contrary, after the strong-wind event, the surface density gradient is less marked (Fig. ). In our
case study the frontal dynamics clearly appear to be an air–sea coupling
process. However, the impact of this coupling depends directly on the wind
regime, as suggested by . During low-wind periods, the NBF
affects the diabatic buoyancy flux and the interactions between the thermal
front and the Ekman pumping generated by wind stress and can also impact the
Ekman buoyancy flux. This interaction is, therefore, air–sea thermal and
dynamic coupling. During strong-wind periods, the NBF position and dynamics
clearly affect the diabatic buoyancy flux. They also affect the wind stress,
but only weakly. There is, therefore, mainly air–sea thermal coupling.
The diabatic buoyancy flux and the adiabatic Ekman buoyancy flux are relevant
parameters to quantify the thermal and the dynamic air–sea coupling effects
on the evolution of stratification. However, this study also shows that advective processes have an important effect on the stratification evolution.
In future work, the various sources of advection could be distinguished in
the stratification budget to better understand the coupling effects,
particularly the Ekman pumping. The interaction between inertial waves and
front could also be studied in detail. For example, show
that the Ekman buoyancy flux is more efficient when the inertial waves reduce
the stratification.
This case study of the North Balearic Front highlights the major role of
Ekman buoyancy fluxes in frontal dynamics and stratification evolution. The
Ekman buoyancy fluxes could significantly influence the open-ocean deep
convection that occurs in the centre of the cyclonic gyre of the Gulf of Lion
during winter as shown by , who highlight a large impact
along the eastern and northern branches of the Northern Current (NC).
suggest that the autumnal NBF dynamics could play a major role in the
preconditioning of the deep convection. Our study shows that this process is
principally driven by Ekman buoyancy fluxes. In future work, the impact of
the Ekman buoyancy flux, all around the cyclonic gyre circulation, could be
quantified during the different phases of the open-ocean deep convection.
Data availability
The Argo data were collected and made freely available by the
CORIOLIS project (http://doi.org/10.17882/42182).
Author contributions
The work was carried out by LS as a PhD candidate under the supervision
of CE, PM and ER. All authors worked on the paper.
Competing interests
The authors declare that they have no conflict of
interest.
Acknowledgements
This work is a contribution to the MISTRALS/HyMeX programme through the
ASICS-MED (ANR-12-BS06-0003) project funded by the French National Agency for
Research (ANR). Data were obtained from the HyMeX programme, sponsored by
grants from MISTRALS/HyMeX and Météo-France. The authors acknowledge the
international Argo programme, the LEFE/GMMC programme and the French NAOS
project for supporting the deployment of profilers. Argo and CTD data were
collected and made freely available by the CORIOLIS project
(http://www.coriolis.eu.org, last access: 8 February 2019) and programmes that contribute to it. We
acknowledge the crews of R/V Suroit and Tethys II and the scientists involved
in the different cruises mentioned in this paper. Numerical simulations were
performed using HPC resources of CALMIP (CALcul en MIdi-Pyrénées,
projects 1247, 09115 and 1325) and GENCI (Grand Equipement National de Calcul
Intensif, project 010569).
Edited by: Simon Josey
Reviewed by: two anonymous referees
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