The Mediterranean Outflow Water (MOW) is a dense water
mass originated in the Strait of Gibraltar. Downstream of the Gulf of Cádiz,
the MOW forms a reservoir region west of the Iberian continental slopes at a
buoyant depth of approximately 1000 m. This region plays a key role as the
main centre where the MOW is mixed and distributed into the North Atlantic.
The seafloor in this area is characterized by the presence of a complex
bathymetry with three abyssal plains separated by mountain chains. Although
the topographic features do not reach the surface, they influence ocean
flows at intermediate and deep ocean layers, conditioning the distribution
and circulation of MOW.
The Copernicus Marine Environmental Monitoring Service (CMEMS) Iberian–Biscay–Ireland (IBI) ocean reanalysis is used to provide a detailed view of the
circulation and mixing processes of MOW near the Iberian and African
continental slopes. This work emphasizes the relevance of the complex
bathymetric features defining the circulation processes of MOW in this
region. The high resolution of the IBI reanalysis allows us to make a
description of the mesoscale features forced by the topography. The
temperature, salinity, velocity, transport, and vorticity fields are
analysed to understand the circulation patterns of MOW. The high-resolution
circulation patterns reveal that Horseshoe Basin and the continental
slope near Cape Ghir (a.k.a. Cap Rhir or Cabo de Aguer) are key areas controlling the mixing processes of MOW
with the surrounding water masses, mainly North Atlantic Central Water (NACW)
and Antarctic Intermediate Water (AAIW). The water mass
variability is also analysed by means of composite analysis. Results
indicate the existence of a variability in the MOW tongue which retracts and
expands westwards in opposition to the movement of the underlying North
Atlantic Deep Water.
Introduction
The Mediterranean Outflow Water (MOW) is a saline and warm water mass
principally occupying the intermediate depths of the eastern North Atlantic. It
is generated by the outflow of subsurface Mediterranean Sea Water (MSW)
through the Strait of Gibraltar. After exiting the Strait of Gibraltar, the
denser MSW cascades along the slope in the Gulf of Cádiz and progressively
entrains the ambient North Atlantic Central Water (NACW). Its downward flow
is disrupted by several transverse submarine valleys splitting up the stream
into two different branches of different densities. Both branches flow along
different erosive channels but finally converge near Cape St. Vincent
forming, on the way, two differenced layers of the same water mass (Iorga and
Lozier, 1999a, b; Gasser et al., 2017). MOW exits the Gulf of Cádiz
at a buoyant depth around 1000 m and afterward it turns northwards following
the western margin of the Iberian Peninsula. It enters the Tagus Basin where
it turns anticyclonically forming a reservoir of this water mass (Daniault
et al., 1994). From this reservoir zone, the MOW spreads into the North
Atlantic, following two main advective pathways: the westward branch towards
the central Atlantic and the poleward branch that, driven by the Iberian
Poleward Current, follows the western European continental slope. On its way
north, MOW enters the Bay of Biscay and continues further north towards
Porcupine Bank and Rockall Trough at 53∘ N (Reid, 1979,
1994; van Aken and Becker, 1996; Iorga and Lozier, 1999a, b; Bower et
al., 2002). Downstream of the poleward pathway, diapycnal mixing occurs with
the underlying Labrador Sea Water core (LSW). It is one of the recognized
ways of salt transport into the inner North Atlantic (Talley and McCartney,
1982; van Aken, 2000). The influence of
MOW on the Nordic seas has been discussed in several studies (Reid, 1979, 1994; Bower et al., 2002; Iorga
and Lozier, 1999a, b; New et al., 2001; McCartney and Mauritzen, 2001);
however, later studies suggest that westward shifts of the subpolar front
controls the input of Atlantic water through Rockall Trough (Lozier and Stewart, 2008).
In addition to these two main paths, the MOW spreads from the reservoir in
south and southwest directions under the influence of other processes and
interacts with surrounding water masses. Thus, another core of MOW spreads
southwestwards largely supported by the motion of Mediterranean Water
lenses, so-called meddies (Armi and Zenk, 1984; Armi et al., 1989; Bower et
al., 1997; van Aken, 2000). The saline signal of MOW extends southward,
reaching the Canary Islands where the water mass collides with a diluted form of
Antarctic Intermediate Water (AAIW) (Machín et al., 2006;
Machín and Pelegí, 2009).
Several works have focused on the description of MOW in the Gulf of Cádiz,
some of them based on observational in situ data (García-Lafuente et
al., 2006; Machín et al., 2006; Machín and Pelegí, 2009) and others based on model products
(Artale et al., 2006; Gasser et al., 2017; Izquierdo and Mikolajewicz,
2019). The consistency between both approaches has improved the knowledge of
the oceanographic processes taking place in this area. However, the MOW flow
beyond the Gulf of Cádiz has not been investigated as extensively and the
scientific literature is rather limited. Most of the studies in this area
are based on compiling data sets of hydrographic stations or climatological
data (e.g. Mazé et al., 1997; Iorga and Lozier, 1999a; van Aken, 2000;
Carracedo et al., 2014). Despite the fact that these methodologies consider a
fair number of stations, they have a limited spatio-temporal coverage,
therefore the results may be influenced by seasonal or mesoscale processes.
On the other hand, the studies based on numerical simulations are usually
focused on larger scales such as the eastern North Atlantic basin scale
(e.g. Iorga and Lozier, 1999b; Bozec et al., 2011) with a resolution too
coarse to provide a proper representation of mesoscale processes at a
regional scale. However, several scientific studies have described the
existence of important mesoscale features in this region, such as the formation of Mediterranean Water lenses (meddies; Armi and Zenk, 1984; Armi et al.,
1989; Bower et al., 1997; Daniault et al., 1994; Sangrá et al., 2009),
development of an anticyclonic gyre centred in the Tagus Abyssal Plain, and
flow along the Gorringe Bank; all highlighting the influence of bathymetry
on the development of mesoscale flows (Zenk and Armi, 1990; Daniault et al., 1994;
Iorga and Lozier, 1999a, b). Therefore, we have a good understanding
of the general circulation of MOW, but there is still a lack of knowledge
about the high-resolution circulation patterns affecting this water mass along its way.
Ocean reanalyses nowadays provide consistent and realistic 3-D gridded ocean
fields over the last decades. The combination of numerical ocean modelling
and data assimilation techniques allows us to get a realistic view of the
circulation patterns including (1) high-enough spatial resolution to
reconstruct the mesoscale features and (2) long temporal coverage allowing us
to compute averaged fields, thus filtering the temporal variability and
retaining stationary features. The present work uses the Copernicus Marine
Environmental Monitoring Service (CMEMS) Iberian–Biscay–Ireland (IBI)
reanalysis (Levier et al., 2014; Sotillo et al., 2015; Aznar et al., 2016;
Amo Baladrón et al., 2018) to make a description of the mean circulation in a
domain encompassing the Tagus, Horseshoe, and Seine abyssal plains. It
emphasizes the mesoscale features together with the general circulation
patterns and their inter-annual variability. The CMEMS IBI reanalysis has
been produced by the CMEMS IBI Monitoring and Forecasting Centre (IBI-MFC)
and, it is one of the multi-year data products used in the Copernicus Ocean
State Report (von Schuckman et al., 2016, 2018,
which aims to provide regular and systematic reference information on the
physical state, variability, and dynamics of the global ocean and European
regional seas. In the framework of its second issue, the IBI-MFC proposed an
ocean monitoring indicator to characterize the inter-annual variability in
MOW in the IBI region (Pascual et al., 2018); to that purpose, an analysis
of the circulation patterns, transports, and temporal variability in MOW was
performed. The present work summarizes part of the results from the
analysis done specifically for the design phase of the CMEMS MOW ocean
indicator. The high resolution of the IBI reanalysis (1/12∘,
∼8 km) allows us to observe several circulation features that
reveal the important role of bathymetry in the spreading of MOW beyond the
Gulf of Cádiz. Additionally, the long temporal coverage of the IBI
reanalysis (1992–2016) permitted the analysis of the mid-term variability in
MOW in the region.
The paper is organized as follow: a preliminary background of the
distribution of MOW near the African and Iberian continental slopes is given
in Sect. 2. Description of the IBI reanalysis and the methods used in the
following sections are presented in Sect. 3. Results are presented and
discussed in Sects. 4 and 5: the analysis of flow, transports, and
hydrographic properties is split into two subsections focused on the Tagus and
Horseshoe basins (Sect. 4.1) and the other on the African slope near Cape
Ghir and Cape Sim (Sect. 4.2); the temporal variability in MOW is studied
in Sect. 5. Finally, main conclusions are summarized in Sect. 6.
Map of the study domain showing the geographical features mentioned in
the text. Red dotted lines denote the sections where the transports have been
computed. They define the three study boxes used in the present work: Gorringe Bank box (GBB), Ampere Bank box (ABB), and Cape Ghir box (CGB) as well as a meridional section in the continental platform: Cape St. Vincent section (CVS).
Background on the MOW regional behaviour
The region of study is bounded by the 39 and 31∘ N parallels, and by the 15∘ W meridian
to the west and the Iberian and northwest African continental shelf edges
to the east (Fig. 1). This domain comprises three abyssal plains oriented
from north to south (Tagus, Horseshoe, and Seine) with depths greater than
5000 m. These three abyssal plains are separated by a number of seamounts
structured into two zonal mountain chains, whose highest peaks are Ampere Bank
and Gorringe Bank. The west margin of the plains is limited by the presence
of the Azores–Portugal Rise. These promontories are quite high obstacles
that clearly separate the three basins, which are delimited by a set of
walls around each plain. Several works on the MOW distribution have
concluded that the complex orography found in this region affects the MOW
flow. The outflow MOW from the Gulf of Cádiz is strongly affected by the
existence of a narrow passage between Gorringe Bank and Cape St. Vincent
(Fig. 1). Zenk and Armi (1990) proposed a schematic flow pattern for the
MOW after its passage through this “gateway”. Later studies have concluded
that despite the fact that most of the MOW flowing out of the Gulf of Cádiz turns
northward through this narrow passage (Daniault et al., 1994), a second
branch of the flow is deflected westwards along the southern flank of the
Gorringe Bank (Iorga and Lozier, 1999a, b).
Another feature commonly described in works is the formation of two gyres.
The first one, a cyclonic gyre, is attached to the Gulf of Cádiz centred at
∼35∘ N, ∼9∘ W with a diameter of ∼250 km
(Iorga and Lozier, 1999a, b), and the second one is an anticyclonic gyre centred in the Tagus
Basin. The latter is the main gyre responsible of the accumulation area known as
MOW reservoir (Iorga and Lozier, 1999a, b). It is worth mentioning
that the definition of the MOW reservoir has been widely used in several
studies; however, the geographic window used to define this area may differ
considerably from the more restrictive definitions (e.g. Iorga and Lozier,
1999a; Pascual et al., 2018) to the more relaxed ones (e.g. Potter and
Lozier, 2004; Bozec et al., 2011).
The hydrographic properties of MOW that reside in the reservoir are
considerably different than the ones of the original MSW generated in the
Strait of Gibraltar. This is due to the intense mixing processes taking
place in the Gulf of Cádiz, which mainly imply the downwelling and
entrainment of NACW (Jia, 2000; Carracedo et al., 2014), as well as of other
ambient water masses. As the reservoir is mostly composed by the mixing of
MSW and NACW, a possible source for the variability in MOW properties in the
reservoir could be the variability in MSW and/or NACW (Artale et al., 2006;
Fucso et al., 2008). However, other studies (Baringer and Price, 1997;
Lozier and Sindlinger, 2009) stated that these two possibilities are too
weak to explain the observed variability in MOW in the reservoir. In a more
recent study, Bozec et al. (2011) concluded that MOW reservoir variability
can be explained by changes in the North Atlantic circulation resulting in a
shift of the preferred MOW pathway, such changes being induced by the
variability of the atmospheric forcing.
The area covered by the Tagus anticyclonic gyre is a key location from where
the MOW is distributed on its two main advective pathways. North of the
Tagus gyre, part of the flow diverges and continues northward beyond the
Extremadura Promontory. The southern boundary of the Tagus gyre is linked
with the westward MOW branch defined by Reid (1994), This branch originates
in the southern flank of Gorringe Bank, and is fed by the splitting of the
flow leaving the Gulf of Cádiz. It travels westwards along the northern wall
of Horseshoe Basin (Iorga and Lozier, 1999a) where it converges with the
southward geostrophic flow coming from the Tagus gyre (Iorga and Lozier, 1999a;
Daniault et al., 1994). The formation of the so-called westward pathway
geographically starts when the flow detaches from the mountain chain and
penetrates westwards into the North Atlantic. However, Iorga and Lozier (1999a, b),
using 80 years of hydrographic data and a diagnostic
model, found that this branch does not count as a significant input of MOW
into the subtropical gyre.
Some evidence suggests that there is some advective southward flow along the
African continental slope. It is originated in the cyclonic gyre centred in
the Gulf of Cádiz, spreads southwards along the African continental slope
(Iorga and Lozier, 1999a), and has been observed at latitudes of
32∘ N (Machín and Pelegrí, 2009). However, later
studies (Izquierdo and Mikolajewicz, 2019) have stated that the tidal
forcing play a major role limiting the southward advective transport of MOW
and contributing to the advection of MOW west from the Gulf of Cádiz. The
general scientific consensus on the spread of MOW south of the Gulf of Cádiz
points out that it is largely supported by the development of meddies
created southwest of the MOW reservoir and near the African continental
slope (Zenk et al., 1992; Sangrá et al., 2009). As studies of the
penetration of the westward branch into the North Atlantic does not show a
clear advective transport of MOW into the subtropical gyre, the formation of
meddies is considered the main cause of the observed large-scale westward
penetration of Mediterranean salt (Arhan and King, 1995).
The southern boundary of the MOW is affected by the interaction with the
overlying AAIW, explaining the deeper salinity signature of MOW at these
latitudes (∼1200 m) due to the presence of the AAIW core at
about 800 m depth. According to Machín and Pelegrí (2009), the
northern limit of AAIW is located around 32∘ N near the
African continental slope; however, their presence northward of the Canary
Islands is subjected to seasonal variability. The mixing processes between
both water masses have been studied by van Aken (2000); they conclude that
the core of AAIW appears to contribute to the formation of MOW since it is
entrained into the overflow near Gibraltar. Such entrainment gives rise to
an enhanced concentration of the nutrients in the Mediterranean originated
water in the North Atlantic.
Several studies have focused on the temporal variability in the MOW and have
concluded that their variability is within a wide range of timescales.
Prieto et al. (2013) analysed a dataset of semi-annual hydrographic stations
located on the Iberian platform. Their findings point out that despite a
larger inter-annual variability, the seasonal signature represents 20 % of
the inter-annual variability. This variability has been found along the
continental slopes affected by MOW, and is a consequence of the variability
of the slope current responsible of the northwards transport of MOW along
the European continental slopes. In summertime, when the narrow jet is
trapped along the continental slope, its thermohaline signature is
reinforced near the slope and diminished its presence in the open
ocean. In wintertime, the slope current weakens and can even be reversed in
some areas (Fricourt et al., 2007; Prieto et al., 2013; Pascual et al., 2018).
Variability at longer timescales has also been addressed by several
studies. Potter and Lozier (2004) analysed 40 years of hydrographic data to
calculate temperature and salinity trends of the MOW reservoir. Along this
period, they found positive temperature and salinity trends that lead to a
heat content gain that overpasses the average gain of the North Atlantic
basin over the later half of the 20th century. Leadbetter et al. (2007)
using three repeated sections at 36∘ N in the North
Atlantic, found an increase in MOW salinity from 1959 to 1981 followed by an
almost compensated decrease in salinity from 1981 until 2005 in the
upper-intermediate layers. According to their results, this change is
controlled by water mass changes along neutral-density surfaces suggesting a
change in the source waters. Bozec et al. (2011) hypothesized that MOW water
mass distribution may be altered by changes in the circulation of the North
Atlantic; to investigate this possible source of variability, they used a set
of model runs forced by either a climatological forcing or an inter-annual
atmospheric forcing. They found a connection between salinity anomalies
along the northern and the westward pathways, thereby concluding that the
observed salinity changes in the MOW reservoir can be explained by
circulation-induced shifts in the salinity field in the eastern North Atlantic basin.
Data and methods
The present study uses the IBI ocean reanalysis delivered by CMEMS (Levier
et al., 2014; Sotillo et al., 2015; Aznar et al., 2016; Amo Baladrón et al.,
2018). The domain covered by the IBI reanalysis is limited by the
26 and 56∘ N parallels, and the
19∘ W and 5∘ E meridians. It provides daily
averages of zonal and meridional velocity components for a period that
ranges from January 1992 to December 2016, with a 1/12∘
horizontal resolution and 75 vertical levels.
The numerical core of the IBI reanalysis is version 3.6 of the eddy-permitting
Nucleus for European Modelling of the Ocean (NEMO) ocean general circulation model (Madec, 2008). This model solves the
three-dimensional finite-difference primitive equations in spherical
coordinates discretized on an Arakawa-C grid. It assumes hydrostatic
equilibrium and the Boussinesq approximation. Vertical mixing is parameterized
according to a k-ε model implemented in the generic form
proposed by Umlauf and Burchard (2003) including mixing induced by surface wave breaking. The bathymetry is derived from GEBCO 08 dataset (Becker et
al., 2016), merged with several local databases.
The IBI run is forced with atmospheric fields from the ECMWF ERA Interim
(Dee et al., 2016). The 10 m wind, surface pressure (added with inverse
barometer approximation), 2 m temperature, and relative humidity are provided at 3 h intervals. On the contrary, precipitation and radiative fluxes are
provided as daily averages from a modified ERA-Interim reanalysis (Sotillo
et al., 2015). CORE empirical bulk formulae (Large and Yeager, 2004) are
used to compute latent sensible heat fluxes, evaporation, and surface stress.
The IBI reanalysis uses lateral open boundary data from the CMEMS global
reanalysis (temperature, salinity, velocities, and sea level) at a resolution
of 0.25∘ (Garric and Parent, 2018). These are complemented
by 11 tidal harmonics built from FES2004 (Lyard et al., 2006) and
TPXO7.1 (Egbert and Erofeeva, 2002). Fresh water river discharge is
implemented as lateral open boundary condition for 33 rivers from
observational and climatological data. Additionally, an extra coastal runoff
rate climatology is used to make the IBI forcing consistent with the ones
imposed in the global system (Maraldi et al., 2013).
The data assimilation scheme applied is the MERCATOR Ocean SAM2 (Lellouche
et al., 2013), established from a singular extended evolutive Kalman filter.
Measurements from CMEMS CORA product (Cabanes et al., 2013; Gatti and
Pouliquen, 2017) database are assimilated as well as high-resolution sea
surface temperature data obtained from analysis of multi-satellite and
Advanced Very High Resolution Radiometer (AVHRR).
products and remote sensing sea level anomalies (SLA) measured by radar
altimeter (Jason-3, Sentinel-3A, HY-2A, Saral/AltiKa, CryoSat-2, Jason-2,
Jason-1, T/P, ENVISAT, GFO, ERS1/2).
This study is based on the regional analysis of the IBI averaged circulation
between 500 and 1500 m for the period 1992–2016. IBI reanalysis provides
ocean fields at 947 and 1045 m depth; therefore, in this work, the 1000 m
depth fields shown are derived from the average of these two IBI levels. The
resulting average circulation field offers information about the main
transports occurring in the region. Since the time average is computed from
daily data at 1/12∘ resolution, the analysed fields include
the net mesoscale transports, filtering its intrinsic temporal variability.
The inter-annual variability in MOW and its associated oceanic patterns is
studied by the analysis of the high/low salinity events in Horseshoe Basin
at 500–1500 m depth. These events are defined as the 10th and 90th percentiles
of the salinity time series. The composites of temperature and
salinity are derived from these events providing an image of the average
ocean state under these conditions.
Part of the analysis is focused over several specific areas (boxes and
sections defined in Fig. 1). These regional domains were selected
when taking into account general circulation patterns. Horseshoe Basin is split
into two different boxes: the northern one (Gorringe Bank box, GBB) comprises
Gorringe Bank and the adjacent western mountain chain, the southern box
(Ampere Bank box, ABB) surrounds the Ampere Bank and the seamounts that form
the southern boundary of Horseshoe Basin. The third box is defined north of
the Canary Islands near to Cape Ghir (Cape Ghir box, CGB). Its limits were
established surrounding the bathymetric promontory located in the African
continental slope between Cape Ghir and Cape Sim. Additionally, a meridional
section south of Cape St. Vincent (Cape St. Vincent section, CVS) has been
defined to compute the net transport and water properties flowing out of the
Gulf of Cádiz. The analysis of water properties, circulation, and transports
is made by averaging over each box (from 500 to 1500 m), as well as along
the sections limiting the boxes and CVS. The volume transports shown are
computed between 500 and 1500 m considering the vertical surface of each
grid cell. The confidence interval of transports has been estimated by
bootstrapping techniques.
MOW regional circulation patternsTagus and Horseshoe basins
Figure 2 shows the mean 1000 m velocity field derived from IBI reanalysis.
MOW transport along the northern slope of Gulf of Cádiz appears as a narrow
stream that flows westward along the slope. The averaged velocity in this
region can reach 0.5–1 m s-1, a value consistent with the current velocities
found in the literature (Gasser et al., 2017; Sánchez-Leal et al.,
2017). South of the outflow stream, results suggest the existence of an
eastward countercurrent that provides the external water input required to
feed the intense entrainment processes generated by MOW in the region (Ambar
and Howe, 1979; Ochoa and Bray, 1991; Baringer and Price, 1997; Jia, 2000).
Velocity field at 1000 m given by the CMEMS IBI reanalysis. For
clarity reasons, grid points with very high relative velocities have been
masked near the southern Iberian platform.
Beyond Cape St. Vincent the flow is strongly affected by the orography.
After the MOW exits the Gulf of Cádiz, the main stream splits when
encountering Gorringe Bank. One part of the flow turns northward through the
narrow passage between Cape St. Vincent and Gorringe Bank and the other part
continues westward south of this promontory. The northward path develops an
intense stream that follows the Portuguese continental slope and enters the
Tagus Basin. After the flow crosses the gateway, it diverges again to form
two separated circulation areas: the bathymetric channelling of water
towards the Tagus Basin that leads to the formation of the widely described
Tagus anticyclonic gyre and a narrow northward flow that, despite the
orographic elevations, goes northward following the margin defined by the
continental slope.
The area formed by the Tagus Basin and Gorringe Bank can be seen as the
centre where the MOW is distributed into the North Atlantic. The narrow
passage between Gorringe Bank and Cape St. Vincent forces the splitting of
the flow exiting the Gulf of Cádiz and promotes the formation of the two
main advective pathways of MOW. The westward MOW pathway begins when the
water masses detach from the continental slope south and north of Gorringe
Bank and travel following the chain of seamounts west of this promontory. On
the southern flank of the mountain range, the flow is forced by the limited
transport through the Gorringe gateway; on the northern flank of Gorringe
Bank, the westward flow follows the southern limit of the Tagus anticyclonic
gyre. The northward branch of MOW commences after the MOW enters the Tagus
Basin. Here the flow splits again into two separated features: the Tagus
anticyclonic gyre, and the poleward slope current which starts near a
promontory in the Portuguese slope (37.8∘ N,
9.4∘ W) and travels northward up to the Extremadura
Promontory. There, the northward flow converges with the northern closure of
the Tagus anticyclonic gyre and continues northward forming the poleward
slope current that will follow the European continental slope up to Porcupine Bank.
Vorticity field at 1000 m derived from IBI velocities. Red/blue
colours denote cyclonic/anticyclonic vorticity, respectively.
The circulation in Horseshoe Basin is mainly characterized by two opposite
zonal currents. The northern half of the basin hosts the westward current
originated in Gorringe Bank, confirming the results suggested by Iorga and
Lozier (1999a); this flow turns cyclonically once it reaches the western
margin of the basin. In the southern half of the basin, the turning flow is
channelized eastward under the influence of the mountain chain of Ampere
Bank. According to our results this flow continues eastward exiting
Horseshoe Basin and returns back to the Gulf of Cádiz, where it provides
part of the water masses that will be entrained by the Mediterranean Water
masses cascading in the Gulf of Cádiz.
The analysis of the mean vorticity field at 1000 m (Fig. 3) provides
information about the main gyres occurring in the area. It is worth
mentioning that this field is affected by the shear of the flow in the
vicinity of the continental slopes. This explains the noisy values detected
near the continental margins. The strong shear between the descending MOW
flow and the surrounding water masses in the northern slopes of the Gulf of
Cádiz generates, in this area, the highest cyclonic vorticity of the figure.
This vorticity value denotes the intense mixing processes taking place in
the area. The bathymetry also has an important influence: high negative values of vorticity are associated with all of the sea
mountains included in the domain, indicating anticyclonic circulation around them, even when the top
of the obstacle is hundreds of metres below the level of 1000 m depth, as
is the case for the sea mount at 36.4∘ N,
13.0∘ W whose summit is at 1893 m depth. Tagus Basin shows
negative values of vorticity mainly related to the presence of the Tagus
anticyclonic gyre. However, results in Horseshoe Basin show the presence of
a generalized cyclonic circulation with two separated centres of vorticity
located at 14 and 12∘ W.
Regarding the cyclonic gyre near the Gulf of Cádiz suggested, by previous
studies (Iorga and Lozier, 1999a, b), the reanalysis does not
provide a clear signal of positive vorticity centred around
35∘ N, 9∘ W. Results in this area mainly show
a zonal westward flow advecting water into the Gulf of Cádiz. In this
area, the water transported eastward diverges taking two directions: part of
the flow turns northward to be reincorporated into the main MOW current in
the Iberian slopes, and a second part of the flow turns anticyclonically
following the African slope towards the Seine Abyssal Plain. The centre of
this circulation pattern appears on the map of vorticity near the slope as
an area of negative values at 34.2∘ N, 9.2∘ W. The splitting of the eastward flow entering the Gulf
of Cádiz into two branches at 9.4∘ W is favoured by the
bathymetric zonal elevation at 35.3∘ N in the Gulf of Cádiz.
Volume transports have been computed across the limits of the two boxes
defined around Gorringe Bank and Ampere Bank (GBB and ABB, respectively). The
analysis of transports through the limits of these boxes provide information
about the water masses entering/leaving Horseshoe Basin as well as the
meridional transport there.
Figure 4 shows the mean transverse velocities and net volume transport in
the sections defined in GBB and ABB. Results show the main input of MOW into
Horseshoe Basin comes through the northern and eastern limits of GBB. The
presence of Gorringe Bank highly influences the MOW flow coming from Cape
St. Vincent. Part of this flow is deflected towards Horseshoe Basin
following the southern flank of Gorringe Bank and enters the basin through
the eastern-GBB section (0.7±0.3 Sv, sverdrup). The input through the northern
limit of GBB is induced by the seamounts at 36.2∘ N, 14.5∘ W which forces a southward transport of water from the
Tagus anticyclonic gyre towards the Horseshoe Basin. The net transport
across sections provides evidence that there is an appreciable flow entering the
basin through the eastern-GBB boundary (0.7±0.3 Sv); however, the flow
through the northern-GBB boundary (1.3±0.3 Sv) is almost twice
higher and is the main source of MOW in Horseshoe Basin.
The combined transport of water into Horseshoe Basin (northern-GBB + eastern-GBB
boundaries) is coherent with the results found in Carracedo et
al. (2104). However, while in this work the transports have been computed
between two fixed levels, in Carracedo et al. (2014) they are computed for
separated water masses. Therefore, the results obtained here are similar to
the transport that Carracedo et al. (2014) labelled as Mediterranean Water
and recirculated Central Water.
Regarding the water outputs in GBB, Fig. 4 shows the presence of a
westward transport that overpasses the seamounts and leaves the box by
crossing the western boundary. This transport of 0.8±0.3 Sv is the
start of the MOW westward pathway described in literature (Iorga and
Lozier, 1999a, b). The latitudes where this westward flow occurs
range from 35.6 to 37.2∘ N.
The net meridional transport between the GBB and ABB is shown in Fig. 4.
The southward transport of water (1.0±0.3 Sv) in Horseshoe Basin is
mainly induced by the cyclonic turning of the westward current previously
described in the basin, in agreement with the results obtained by Iorga and
Lozier (1999a, b). The southward penetration of this water is stopped
by the southern wall of the basin, redirecting this flow eastward. However,
results suggest a net external water input (0.8±0.5 Sv) entering the
basin through the southern boundary of ABB. The two flows then converge and
exit ABB eastwards towards the Gulf of Cádiz (1.8±0.4 Sv). The
θ/S diagrams averaged in GBB and ABB (Fig. 4b) show the
differences in temperature and salinity between the northern and southern
halves of Horseshoe Basin. Such differences mainly affect the
upper-intermediate layers of ABB where a freshening and cooling of waters is
appreciable. The water flow averaged in GBB presents the usual
θ/S profile of MOW with a peak in salinity around 1000 m depth (Leadbetter et al.,
2007). On the contrary, the θ/S profile averaged in ABB shows some
mixing of AAIW and North Atlantic Deep Water (NADW). The salinity peak is
still found at approximately 1000 m depth; however, the influence of AAIW is
proven by the temperature and salinity reduction in the layers above the
salinity peak (from 600 to 1000 m depth) producing a concavity that sharpens
the peak. The influence of NADW is reflected in the freshening and cooling
in deeper layers.
Despite of the net southward transport between GBB and ABB boxes (Fig. 4a),
the averaged velocities across their shared boundary reveal smaller
northward flows advecting water from ABB to GBB; this implies the existence
of meridional mixing processes in Horseshoe Basin. The northward transport
of modified water can be appreciated in the velocity and vorticity fields
shown in Figs. 2 and 3: the general circulation in Horseshoe Basin is
composed by two separated centres of cyclonic vorticity, located at
35.7∘ N, 14.2∘ W and at
35.7∘ N, 12.0∘ W. The circulation around
these two centres explains the north–south flows in Horseshoes Basin. The
θ/S profile averaged at the eastern boundary of ABB (Fig. 4b)
proves that the properties of water exiting the basin towards the Gulf of
Cádiz are the result of the mixing between the southern and northern waters
in Horseshoe Basin. However, the modification of properties mainly affects
the upper-intermediate layers above 1100 m depth.
(a) Transverse velocity (black arrows) and net volume
transport (red arrows) in sections delimiting GBB and ABB. (b) Mean
θ/S diagram averaged in GBB (circles), ABB (squares), and along the
section marking the eastern-ABB boundary (diamonds). Dotted isolines correspond
to potential density anomaly in kilogram per cubic metre (kg m-3). Black dots represent the position
of the source water types AAIW and MOW.
(a) Transverse velocity (black arrows) and net volume transport
(red arrows) in sections delimiting CGB. (b) Mean θ/S diagram
averaged along the sections delimiting the northern-CGB boundary (circles),
southern-CGB boundary (squares), and western-CGB boundary (diamonds). Dotted
isolines correspond to potential density anomaly in kilogram per cubic metre (kg m-3). Black dots
represent the position of the source water types AAIW and MOW.
(a) Anomalies of salinity at 1000 m averaged in the whole
Horseshoe Basin (areas GBB and ABB combined) and in CVS (blue and grey lines,
respectively). Red/green dots depicts the values under/over the
10h/90th percentile of salinity anomaly in Horseshoe Basin. (b)θ/S
diagram in Horseshoe Basin (areas GBB and ABB combined) averaged in the complete
time record (squares), dates of minimum salinity (circles), and dates of maximum
salinity (diamonds). Dotted isolines correspond to potential density anomaly
in kilogram per cubic metre (kg m-3). Black dots represent the position of the source water types
NADW and MOW.
Cape Ghir
The average circulation pattern at 1000 m shown in Fig. 2 reveals the
existence of a zonal current at 31.5∘ N. It is formed along
the African continental slope and confirms the existence of an advective
westward flow that penetrates into the North Atlantic up to
16.2∘ W. This current is associated with an area of cyclonic
vorticity west of Cape Ghir and centred at 30.5∘ N and
11.6∘ W (Fig. 3). The domain CGB, near Cape Ghir, has
been defined to analyse the source of this flow (Fig. 1). Figure 5
displays the velocity and water transport at the boundaries of CGB. It
reveals the entrance of water masses from the north as a weak southward
transport (0.6±0.3 Sv) between 12 and
11∘ W, and from the south as a more intense transport
(2.9±0.2 Sv) close to the continental slope. The exit of water masses
occurs across the western limit of CGB where a westward current of 5–8 cm s-1
(3.3±0.3 Sv) pushes the converging water masses towards the outer
ocean (Fig. 5a). The θ/S diagram averaged in the boundaries of the
domain shows the presence of the salinity peak at 1000 m associated with the
MOW (Fig. 4b). However, its temperature and salinity values are
significantly lower than ones obtained 500 km to the north at Horseshoe Basin,
where the MOW salinity peak ranges between 35.9–36.0 psu and
10.1–10.4 ∘C. The water layers above the salinity peak reveal the
strong influence of the AAIW through the presence of a concavity in the
profile associated with the cooler and fresher AAIW. The result found in the
northern section of CGB is consistent with the observed climatological
θ/S diagram averaged from 10 to 12∘ W and from
31 to 33∘ N by Machín and
Pelegrí (2009). Comparing the θ/S profiles of the northern and
southern boundaries of the CGB, results show a different relative
contribution of MOW and AAIW in each section. While in the southern-GCB
boundary, the θ/S profile shows a dominance of AAIW with reduced
values of salinity and a concavity in the upper-intermediate waters; at the
northern-GCB section, the θ/S diagram shows a greater influence of
MOW with a sharper peak of salinity at 1000 m. According to the
θ/S diagram averaged at the western boundary of the box, the outgoing waters
reveal a profile influenced by both the water masses of the northern and
southern boundaries of GCB. The peak of salinity diminishes and the values
of temperature and salinity in the upper-intermediate layers increase
reducing its concavity. The analysis of the mean velocities shown in
Fig. 2 provides more information about the circulation processes taking place in
this area. South of Cape Ghir, the poleward along-slope current mainly
composed by AAIW follows the continental slope up to the promontory near
Cape Ghir and even further (Machín and Pelegrí, 2009). There, the
promontory and the opposition of the southward flow of MOW forces the
cyclonic turning of AAIW that detaches from the continental slope and starts
an advective westward flow penetrating up to 16.2∘ W. The
mixing, resulting from this confrontation of water masses, leads to the
modification of temperature and salinity seen in the θ/S profile
averaged in western-GCB boundary.
Inter-annual variability
To analyse the inter-annual variability in MOW in Horseshoe Basin, the
anomalies of salinity at 1000 m have been computed and spatially averaged in
the area composed by GBB and ABB. The analysis of the time series of
salinity anomalies reveals a different behaviour of the maximum and minimum
anomalies (Fig. 6). While the maximum values of salinity are spread over time, the minimum anomalies are concentrated around a unique event
occurring during the period 2000–2003. This low-salinity event in the area
is clear since all values under the 10th percentile of the time series are
found within these years. This event is consistent with the results found by
Bozec et al. (2011), who report a slight retreat of the inner salinity
contours towards the east, especially after 2000.
As stated in Sect. 2, previous works on the variability in hydrographic
properties in the MOW reservoir have concluded that changes in MOW properties
are not dominated by changes in MSW properties (Lozier and Sindlinger, 2009;
Bozec et al., 2011). This finding is confirmed by the results of this work in
Fig. 6a, which show salinity anomalies averaged in Horseshoe Basin and CVS.
The correlation between both time series is not significant, thus denying
direct-linear relationship between salinity anomalies in these areas. This
correlation remains barely significant even if the time series are lagged:
the maximum significant cross-correlation (0.22) is obtained when the time
series averaged in CVS is retarded 2.5 years (Fig. 7).
Cross-correlation coefficient of salinity anomalies averaged in
Horseshoe Basin (areas GBB and ABB combined) and CVS. Lag corresponds to the
displacement of the CVS time series respect to the Horseshoe Basin time series.
Red segments denote statistically significant correlations with significance
level of 99 %.
The dates of the minimum and maximum anomalies of salinity defined
respectively as the values under/over the 10th/90th percentiles have been
used to derive the averaged fields on these dates. The resulting composites,
associated with high/low salinity situations have been used to analyse (1) the
θ/S diagram averaged in the Horseshoe Basin (Fig. 6b), (2) the
meridional section of salinity at 36∘ N coinciding with the
limit between GBB and ABB (Fig. 8), and (3) the salinity distribution at
the maximum salinity depth (Fig. 9). Since both composites are derived
from the fields averaged at dates where the anomaly at 1000 m depth is
remarkably high or low, they represent the ocean state associated with these
particular salinity conditions. Moreover, as every minimum value selected is
related to the period 2000–2003, the derived minimum composite mainly reflects the ocean state under this specific low-salinity event.
Section of salinity at 36∘ N (coinciding with the limit
between GBB and ABB) of the composites resulting from the averaged fields
defined from the 10th (a) and 90th (b) percentiles of salinity
anomalies in Horseshoes Basin.
(a, b) Maximum salinity depth (shaded colours) between
500 and 1500 m and salinity (isolines) at the maximum salinity depth. Maps
obtained from the composites of minimum (a) and maximum (b)
salinity defined by 10th and 90th percentiles of salinity anomalies in Horseshoe Basin, respectively. (c) Difference in salinity between (a)
and (b) and contour lines represent bathymetry.
The θ/S composites shown in Fig. 6 reveal that the variability in
the Horseshoe Basin mainly affects the layers below the salinity peak. Under
conditions of the minimum salinity event, the water masses below 1000 m
depth suffer a decrease of approximately 0.25 psu. Temperature is also
affected during this event, showing a cooling of about
∼0.6∘C. The structure of the water column is also modified,
the salinity peak associated with the MOW core being pushed from the usual
1000 m up to 800 m depth.
Based on three repeated sections at 36∘ N, Leadbetter et al. (2007)
observed inter-annual variability in intermediate water masses between
10 and 20∘ W. This variability mainly
affected the layers above the salinity peak, whereas variability in lower
levels were smaller. However, our results, based on the use of a high
resolution regional reanalysis, suggest an inconsistency since the
θ/S variability in this area mainly affects the layers below the MOW
salinity peak. As presented in previous sections, variability above the
salinity peak can be rather attributed to latitudinal variations, as the
differences between θ/S diagrams averaged at GBB and in ABB
(∼100 km. southward) are mainly found in the upper levels
(above the MOW salinity peak). On the contrary, according to the present
analysis, the main temporal variability in the Horseshoe Basin is found in
layers below the MOW tongue. The composite sections of salinity at
36∘ N shown in Fig. 8 depict the vertical structure of the
water column under conditions of maximum and minimum salinity anomaly. Under
conditions of the minimum salinity event (Fig. 8a), results suggest a
westward propagation of the underlying NADW that leads to a retreat of the
MOW tongue together with an upward displacement of the water mass core. This
process also implies a general freshening and cooling of the whole water
column. The accumulation of low-salinity waters west of 15∘ W
and below 1100 m supports this hypothesis and suggests that penetration of
the NADW into Horseshoe Basin is limited by the Azores–Portugal Rise. The
perturbation in isohalines can be appreciated up to 400 m above the top of
the seamount. This result agrees with Bozec et al. (2011) who reported a
similar behaviour of the boundary between MOW and the underlying LSW in the
central Atlantic.
Figure 9 analyses the maximum of salinity between 500 and 1500 m depth for
the composites of maximum and minimum salinity at Horseshoe Basin. As seen
previously, during the high-salinity situations, the maximum of salinity
appears deeper than during the low-salinity situations. Additionally, the
overall salinity field becomes saltier, inducing a westward displacement of
isohalines. The bigger salinity differences are found at the northwest limit
of the Horseshoe Basin, near the Azores–Portugal Rise. On the contrary, in
the vicinity of the continental platform, the salinity differences are
smaller, or even negative in the Gulf of Cádiz and surroundings of Cape
St. Vincent. This result suggests a negatively correlated behaviour between the salinity fields of Horseshoe Basin and Cape St. Vincent and Gulf of Cádiz.
Conclusions
In the present work, an analysis of the spreading processes of MOW in the
northeast Atlantic has been performed through the use of a high-resolution
ocean reanalysis: the CMEMS IBI regional reanalysis. The ocean properties
and flows are analysed at intermediate depths (500–1500 m depth) in the
Tagus, Horseshoe, and Seine basins. These basins are adjacent to the Gulf of
Cádiz and they form the main area where the MOW accumulates before spreading
into the North Atlantic. In this work we analyse the tongue of MOW in the
reservoir area, the influence of bathymetry over the spreading of MOW, and
its interactions with the surrounding water masses (NADW and AAIW). This
work has also been conducted by comparing the IBI reanalysis with previous
works on the characterization of MOW in the North Atlantic. The high
agreement of results with the known features of MOW suggests a proper
reproduction of the dynamic features of the intermediate waters in the
region. Moreover, the high resolution of the IBI reanalysis product allows
us to describe new mesoscale features, previously not reported in literature.
One of the main contributions of this work results from the updated
description of circulation patterns of MOW in the east North Atlantic. A
coarse description of the circulation patterns in this area was reported by
previous studies (Daniault et al., 1994; Mazé et al., 1997; Iorga and
Lozier, 1999a, b; van Aken, 2000); however, the high resolution of the IBI
reanalysis allows for a more detailed description of the circulation in the
region. Figure 10 summarizes the circulation patterns described by the
abovementioned works (represented by black arrows) and the updated scheme resulting from
the IBI reanalysis data (red arrows). Once the recently
formed MOW overpasses Cape St. Vincent, the presence of Gorringe Bank splits
the flow into two branches: one of them enters Tagus Basin, describing an
anticyclonic gyre, whereas the other branch of MOW flows westward along the
northern boundary of Horseshoe Basin. The bathymetry also forces the
cyclonic circulation in Horseshoe Basin, the westward flow indicated in
Gorringe Bank turns cyclonically when it encounters the seamounts at the
western boundary of the basin (the Azores–Portugal Rise). Thereafter, water
masses recirculate eastwards towards the Gulf of Cádiz. Seamounts of the
Ampere mountain chain lead to the formation of anticyclonic vorticity in the
proximity of these obstacles, and circulation around these centres favours
active transports of southern water into Horseshoes Basin. The external
water entrained is mainly composed by a diluted form of AAIW in the
upper-intermediate layers and NADW at depths below 1000 m. Thereby, the MOW
water masses that recirculates into the Gulf of Cádiz are previously
modified in the Horseshoe Basin region through mixing with AAIW and NADW. This
process could explain the presence of an enhanced concentration of nutrients
in the MOW, as reported by van Aken (2000).
Schematic representation of the overall Mediterranean Outflow Water
pathways in the eastern North Atlantic. Black arrows show the current scientific
consensus according to Iorga and Lozier (1999a) and Carracedo et al. (2014).
Red arrows summarize other MOW features described in the work. Results based on an analysis of the CMEMS IBI reanalysis (25 years record and
1/12∘ resolution).
The interaction between MOW and the deeper AAIW near the African continental
slope is highly influenced by the bathymetric promontory between Cape Ghir
and Cape Sim. The converging flows of MOW, which travel southward along the
African continental slope, and of AAIW, which enters the basin northwards
through a narrow gateway between Fuerteventura and the African Shelf,
collide in this area and produce a zonal transport of mixed MOW–AAIW waters
towards the inner ocean. The potential role of this branch as an advective
pathway for MOW into the subtropical gyre and its relationships with the
“Madeira Eddy Corridor”, described by Sangrá et al. (2009), will be analysed in future works.
The analysis of the circulation patterns in the region has highlighted the
role of the bathymetry as a key factor determining the spreading and mixing
patterns of MOW in the region. The presence of three abyssal plains
separated by seamount chains implies an orographic complexity that highly
influences the processes in the area. This work has also reported some cases
where the bathymetric features can modify the flow hundreds of metres above
the obstacles, some topographic features moreover coinciding with the
observed areas of meddy formation as described by previous authors
(Sangrá et al., 2009) and suggesting some influence of bathymetric
obstacles on the meddy formation. Therefore, in other to obtain a realistic
representation of the ocean, the influence of the high-resolution bathymetry
must be considered by modelling studies in this region and depths.
The analysis of the CMEMS IBI currents and transports together with the
derived hydrographic patterns has allowed for the reconstruction of the dynamic
variability in Horseshoe Basin. The composite analysis of temperature and
salinity in the region leads to the conclusion that the main source of
inter-annual variability in the Horseshoe Basin comes from the deeper layers of
MOW. Our work has shown that the boundary between the MOW and the underlying
NADW is subject to inter-annual variability. As far as observed within the
limited temporal extension of the CMEMS IBI reanalysis (25 years), the
presence of the MOW tongue in Horseshoe Basin seems to be the normal
situation. However, the reanalysis reveals the existence of a remarkable
event (2000–2004) where the NADW advances into Horseshoe Basin. Under these
specific conditions, the MOW core retreats eastwards, diminishing the
salinity in the whole water column of Horseshoe Basin. The analysis of the
maximum vertical salinity of composites has led us to conclude that the North
Atlantic circulation influences the water properties of the Gulf of Cádiz
and Cape St. Vincent. Thus, the advance of NADW implies the accumulation of
MOW near its source. Therefore, the low-salinity event in Horseshoe
Basin is associated with an increase in salinity in the intermediate layers of the Gulf of Cádiz and Cape St. Vincent.
Data availability
All data used in this study can be downloaded from the
Copernicus Marine Environment Monitoring Service (2019) web page
http://marine.copernicus.eu/services-portfolio/access-to-products/?option=com_csw&view=details&product_id=IBI_REANALYSIS_PHYS_005_002.
Author contributions
All the authors have contributed as follows. AdPC: scientific expertise,
technical support, and text drafting. MGS: scientific expertise and text
drafting. BL: scientific expertise and technical support. RA: scientific
expertise and text drafting. PL: scientific expertise and technical support.
AAV: technical support. EAF: scientific expertise and text drafting.
Competing interests
The authors declare that they have no conflict of
interest.
Special issue statement
This article is part of the special issue “The Copernicus Marine
Environment Monitoring Service (CMEMS): scientific advances”. It is not associated with a conference.
Acknowledgements
The authors thank to Copernicus Marine Environment Monitoring Service for
providing the data. Additionally, the helpful comments of Karen Guihou and
the other referees are gratefully acknowledged.
Review statement
This paper was edited by Marina Tonani and reviewed by two
anonymous referees.
References
Ambar, I. and Howe, M. R.: Observations of the Mediterranean outflow. II. The
deep circulation in the vicinity of the Gulf of Cadiz, Deep-Sea Res., 26A, 555–568, 1979.Amo Baladrón, A., Levier, B., and Sotillo, M. G.: Product User Manual for
Atlantic-Iberian Biscay Irish-Ocean Physics Reanalysis Product: IBI_REANALYSIS_PHYS_005_002,
Copernicus Marine Environment Monitoring Service, available at:
http://cmems-resources.cls.fr/documents/PUM/CMEMS-IBI-PUM-005-002.pdf (17 May 2019), 2018.Arhan, M. and King, B.: Lateral mixing of the Mediterranean water in the eastern
North Atlantic, J. Mar. Res., 53, 865–895, 10.1357/0022240953212990, 1995.
Armi, L. and Zenk, W.: Large lenses of highly saline Mediterranean water, J.
Phys. Oceanogr., 14, 1560–1576, 1984.
Armi, L., Hebert, D., Oakey, N., Price, J., Richardson, P., Rossby, H., and
Ruddinck, B.: Two years in the life of a Mediter- ranean salt lens, J. Phys.
Oceanogr., 19, 354–383, 1989.
Artale, V., Calmanti, S., Malanotte-Rizzoli, P., Pisacane, G., Rupolo, V., and
Tsimplis, M.: The Atlantic and Mediterranean Sea as connected systems, in:
Mediterranean Climate Variability, edited by: Lionello, P., Malanotte-Rizzoli,
P., and Boscolo, R., Elsevier, Oxford, UK, 283–322, 2006.Aznar, R., Sotillo, M. G., Cailleau, S., Lorente, P., Levier, B.,
Amo-Baladrón, A., Reffray, G., and Álvarez-Fanjul, E.: Strengths and
weaknesses of the CMEMS forecasted and reanalyzed solutions for the
Iberia-Biscay-Ireland (IBI) waters, J. Mar. Syst., 159, 1–14, 10.1016/j.jmarsys.2016.02.007, 2016.
Baringer, M. O. and Price, J. F.: Mixing and spreading of the Mediterranean
outflow, J. Phys. Oceanogr., 27, 1654–1677, 1997.Becker, J. J., Sandwell, D. T., Smith, W. H. F., Braud, J., Binder, B., Depner,
J., Fabre, D., Factor, J., Ingalls, S., Kim, S.-H., Ladner, R., Marks, K.,
Nelson, S., Pharaoh, A., Trimmer, R., von Rosenberg, J., Wallace, G., and
Weatherall, P.: Global bathymetry and elevation data at 30 arc seconds
resolution: SRTM30_PLUS, Mar. Geod., 32, 355–371, 10.1080/01490410903297766, 2016.
Bower, A. S., Armi, L., and Ambar, I.: Lagrangian observations of meddy formation
during A Mediterranean Undercurrent Seeding Experiment, J. Phys. Oceanogr.,
27, 2545–2575, 1997.Bower, A. S., Lecann, B., Rossby, T., Zenk, W., Gould, J., Speer, K., Richardson,
P. L., Prater, M. D., and Zhang, H.-M.: Directly measured mid-depth circulation
in the northeastern North Atlantic Ocean, Nature, 419, 603–607, 10.1038/nature01078, 2002.Bozec, A., Lozier, M. S., Chasignet, E. P., and Halliwel, G. R.: On the
variability of the Mediterranean Outflow Water in the North Atlantic from 1948
to 2006, J. Geophys. Res.-Oceans, 116, C09033, 10.1029/2011JC007191, 2011.Cabanes, C., Grouazel, A., von Schuckmann, K., Hamon, M., Turpin, V., Coatanoan,
C., Paris, F., Guinehut, S., Boone, C., Ferry, N., de Boyer Montegut, C.,
Carval, T., Reverdin, G., Pouliquen, S., and Le Traon, P.-Y.: The CORA dataset:
validation and diagnostics of in-situ ocean temperature and salinity measurements,
Ocean Sci., 9, 1–18, 10.5194/os-9-1-2013, 2013.Carracedo, L., Gilcoto, M., Mercier, H., and Pérez, F.: Seasonal dynamics
in the Azores-Gibraltar Strait region: A climatologically-based study, Prog.
Oceanogr., 122, 116–130, 10.1016/j.pocean.2013.12.005, 2014.Copernicus Marine Environment Monitoring Service: Atlantic-Iberian Biscay
Irish-Ocean Physics Reanalysis, available at:
http://marine.copernicus.eu/services-portfolio/access-to-products/?option=com_csw&view=details&product_id=IBI_REANALYSIS_PHYS_005_002,
last access: 21 May 2019.Daniault, N., Mazé, J. P., and Arhan, M.: Circulation and mixing of
Mediterranean water west of the Iberian Peninsula, Deep-Sea Res., 41, 11–12,
10.1016/0967-0637(94)90068-X, 1994.Dee, D. P., Uppala, S. M., Simmons, A. J., Berrisford, P., Poli, P., Kobayashi,
S., Andrae, U., Balmaseda, M. A., Balsamo, G., Bauer, P., Bechtold, P., Beljaars,
A. C. M., van de Berg, L., Bidlot, J. R., Bormann, N., Delsol, C., Dragani, R.,
Fuentes, M., Geer, A. J., Haimberger, L., Healy, S. B., Hersbach, H., Hólm,
E. V., Isaksen, L., Kallberg, P., Köhler, M., Matricardi, M., McNally, A.
P., Monge-Sanz, B. M., Morcrette, J. J., Park, B. K., Peubey, C., de Rosnay, P.,
Tavolato, C., Thépaut, J. N., and Vitart, F.: The ERA-interim reanalysis:
configuration and performance of the data assimilation system, Q. J. Roy.
Meteorol. Soc., 137, 553–597, 10.1002/qj.828, 2016.Egbert, G. D. and Erofeeva, S. Y.: Efficient inverse modeling of Barotropic
Ocean tides, J. Atmos. Ocean. Tech., 19, 183–204, 10.1175/1520-0426(2002)019b0183:EIMOBON2.0.CO;2, 2002.Friocourt, Y., Levier, B., Speich, S., Blanke, B., and Drijfhout, S. S.: A
regional numerical ocean model of the circulation in the Bay of Biscay, J.
Geophys. Res.-Oceans, 112, 1–19, 10.1029/2006JC003935, 2007.Fucso, G., Artale, V., and Cotroneo, Y.: Thermohaline variability of
Mediterranean Water in the Gulf of Cadiz over the last decades (1948–1999),
Deep-Sea Res. Pt. I, 55, 1624–1638, 10.1016/j.dsr.2008.07.009, 2008.García-Lafuente, J., Delgado, J., Criado-Aldeanueva, F., Bruno, M.,
del Río, J., and Miguel Vargas, J.: Water mass circulation on the
continental shelf of the Gulf of Cádiz, Deep-Sea Res. Pt. II, 53, 1182–1197,
10.1016/j.dsr2.2006.04.011, 2006.Garric, G. and Parent, L.: Product User Manual for Global Ocean Reanalysis
Product: GLOBAL-REANALYSIS-PHY-001-025, Copernicus Marine Environment
Monitoring Service, available at: http://cmems-resources.cls.fr/documents/PUM/CMEMS-GLO-PUM-001-025.pdf (last access: 17 May 2019), 2018.Gasser, M. Pelegrí, J. L., Emelianov, M., Bruno, M., Grácia, E., Pastor,
M., Peters, H., Rodríguez-Santana, A., Salvador, J., Sánchez-Leal, R.
L.: Tracking the Mediterranean outflow in the Gulf of Cadiz, Prog. Oceanogr.,
157, 47–71, 10.1016/j.pocean.2017.05.015, 2017.Gatti, J. and Pouliquen, S.: Product user manual for near real time and delayed
mode objective analysis products INSITU_GLO_TS_OA_REP_OBSERVATIONS_013_002_ab
Period covered: 1990–2015, Copernicus Marine Environment Monitoring Service,
available at: http://cmems-resources.cls.fr/documents/PUM/CMEMS-INS-PUM-013-002-ab.pdf (17 May 2019), 2017.
Iorga, M. C. and Lozier, M. S.: Signatures of the Mediterranean outflow from a
North Atlantic climatology 1. Salinity and density fields, J. Geophys. Res.,
104, 25985–26009, 1999a.
Iorga, M. C. and Lozier, M. S.: Signatures of the Mediterranean outflow from a
North Atlantic climatology 2. Diagnostic velocity fields, J. Geophys. Res.,
104, 26011–26029, 1999b.Izquierdo, A. and Mikolajewicz U.: The role of tides in the spreading of
Mediterranean Outflow Waters along the Southwestern Iberian Margin, Ocean Model.,
133, 27–43, 10.1016/j.ocemod.2018.08.003, 2019.Jia, Y.: Formation of an Azores Current Due to Mediterranean Overflow in a
Modeling Study of the North Atlantic, J. Phys. Oceanogr., 30, 2342–2358, 10.1175/1520-0485(2000)030<2342:FOAACD>2.0.CO;2, 2000.Large, W. G. and Yeager, S. G.: Diurnal to Decadal Global Forcing for Ocean
and Sea-Ice Models: the Data Sets and Flux Climatologies, NCAR technical
note, NCAR/TN-460CSTR, National Center for Atmospheric Research, Boulder,
Colorado, 10.5065/D6KK98Q6, 2004.Leadbetter, S. J., Williams, R. G., McDonagh, E. L., and King, B. A.: A twenty
year reversal in water mass trends in the subtropical North Atlantic, Geophys.
Res. Lett., 34, 1–6, 10.1029/2007GL029957, 2007.Lellouche, J.-M., Le Galloudec, O., Drévillon, M., Régnier, C., Greiner,
E., Garric, G., Ferry, N., Desportes, C., Testut, C.-E., Bricaud, C.,
Bourdallé-Badie, R., Tranchant, B., Benkiran, M., Drillet, Y., Daudin, A.,
and De Nicola, C.: Evaluation of global monitoring and forecasting systems at
Mercator Océan, Ocean Sci., 9, 57–81, 10.5194/os-9-57-2013, 2013.
Levier, B., Benkiran, M., Reffray, G., and Sotillo, M. G.: IBIRYS: A Regional
High Resolution Reanalysis (Physical and Biogeochemical) over the European
North East Shelf, EGU General Assembly, id.14014, 2014.Lozier, M. S. and Stewart N. M.: On the temporally varying penetration of
Mediterranean overflow waters and eastward penetration of Labrador Sea Water,
J. Phys. Oceanogr., 38, 2097–2103, 10.1175/2008JPO3908.1, 2008.Lozier, M. S. and Sindlinger, L.: On the Source of Mediterranean Overflow Water
Property Changes, J. Phys. Oceanogr., 39, 1800–1817, 10.1175/2009JPO4109.1, 2009,Lyard, F., Lefevre, F., Letellier, T., and Francis, O.: Modelling the global
ocean tides: modern insights from FES2004, Ocean Dynam., 56, 394–415,
10.1007/s10236-006-0086-x, 2006.
Machín, F. and Pelegrí, J. L.: Northward Penetration of Antarctic
Intermediate Water off Northwest Africa, J. Phys. Oceanogr., 39, 512–535, 2009.Machín, F., Pelegrí, J. L., Marrero-Díaz, A., Laiz, I., and
Ratsimandresy, A. W.: Near-surface circulation in the southern Gulf of Cádiz,
Deep-Sea Res. Pt. II, 53, 1161–1181, 10.1016/j.dsr2.2006.04.001, 2006.
Madec, G.: NEMO Ocean General Circulation Model, Reference Manual, Internal
Report, LODYC/IPSL, Paris, 2008.Maraldi, C., Chanut, J., Levier, B., Reffray, G., Ayoub, N., De Mey, P., Lyard,
F., Cailleau, S., Drévillon, M., Fanjul, E. A., Sotillo, M. G., Marsaleix,
P., and the Mercator team: NEMO on the shelf: assessment of the
Iberia-Biscay-Ireland configuration, Ocean Sci., 9, 745–771, 10.5194/os-9-745-2013, 2013.Mazé, J. P., Arhan, M., and Mercier, H.: Volume budget of the eastern
boundary layer off the Iberian Peninsula, Deep-Sea Res. Pt. I, 44, 1543–1574,
10.1016/S0967-0637(97)00038-1, 1997.McCartney, M. and Mauritzen, C.: On the origin of the warm inflow to the Nordic
Seas, Prog. Oceanogr., 51, 125–214, 10.1016/S0079-6611(01)00084-2, 2001.New, A. L., Barnard, S., Herrmann, P., and Molines, J.-M.: On the origin and
pathway of saline inflow to the Nordic Seas: Insights from the models, Prog.
Oceanogr., 48, 255–287, 10.1016/S0079-6611(01)00007-6, 2001.
Ochoa, J. and Bray, N. A.: Water mass exchange in the Gulf of Cadiz, Deep-Sea
Res., 38, 5465–5503, 1991.Pascual, Á., Levier, B., and Sotillo, M.: Characterisation of Mediterranean
outflow water in the Iberian-Gulf of Biscay-Ireland region, in: Copernicus
Marine Service Ocean State Report, Issue 2, J. Operat. Oceanogr., 11, s1–s142,
10.1080/1755876X.2018.1489208, 2018.Potter, R. A. and Lozier, M. S.: On the warming and salinification of the
Mediterranean outflow waters in the North Atlantic, Geophys. Res. Lett., 31,
1–4, 10.1029/2003GL018161, 2004.Prieto, E., González-Pola, C., Lavín, A., Sánchez, R. F., and
Ruiz-Villarreal, M.: Seasonality of intermediate waters hydrography west of
the Iberian Peninsula from an 8 yr semiannual time series of an oceanographic
section, Ocean Sci., 9, 411–429, 10.5194/os-9-411-2013, 2013.Reid, J. L.: On the contribution of the Mediterranean Sea outflow to the
Norwegian-Greenland Sea, Deep-Sea Res. Pt. A, 26, 1199–1223, 10.1016/0198-0149(79)90064-5, 1979.Reid, J. L.: On the total geostrophic circulation of the North Atlantic Ocean:
Flow patterns, tracers, and transports, Prog. Oceanogr., 33, 1–92,
10.1016/0079-6611(94)90014-0, 1994.Sánchez-Leal, R. F., Bellanco, M. J., Fernández-Salas, L. M.,
García-Lafuente, J., Gasser-Rubinat, M., González-Pola, C.,
Hernández-Molina, F. J., Pelegrí, J. L., Peliz, A., Relvas, P., Roque,
D., Ruiz-Villarreal, M., Sammartino, S., and Sánchez-Garrido, J. C.: The
Mediterranean Overflow in the Gulf of Cadiz: A rugged journey, Sci. Adv., 3,
1–12, 10.1126/sciadv.aao0609, 2017.Sangrà, P., Pascual, A., Rodríguez-Santana, Á., Machín, F.,
Mason, E., McWilliams, J. C., Pelegrí, J. L., Dong, C., Rubio, A.,
Arístegui, J., Marrero-Díaz, Á., Hernández-Guerra, A.,
Martínez-Marrero, A., and Auladell, M.: The Canary Eddy Corridor: A major
pathway for long-lived eddies in the subtropical North Atlantic, Deep-Sea Res.
Pt. I, 56, 2100–2114, 10.1016/j.dsr.2009.08.008, 2009.
Sotillo, M. G., Cailleau, S., Lorente, P., Levier, B., Aznar, R., Reffray, G.,
Amo-Baladrón, A., Chanut, J., Benkiran, M., and Álvarez Fanjul, E.:
The MyOcean IBI Ocean forecast and reanalysis systems: operational products
and roadmap to the future Copernicus Service, J. Operat. Oceanogr., 8, 63–79,
10.1080/1755876X.2015.1014663, 2015.Talley, L. D. and McCartney, M. S.: Distribution and Circulation of Labrador
Sea Water, J. Phys. Oceanogr., 12, 1189–1205, 10.1175/1520-0485(1982)012<1189:DACOLS>2.0.CO;2, 1982.Umlauf, L. and Burchard, H.: A generic length-scale equation for geophysical
turbulence models, J. Mar. Res., 61, 235–265, 10.1357/002224003322005087, 2003.van Aken, H. M.: The hydrography of the mid-latitude Northeast Atlantic Ocean
II: The intermediate water masses, Deep-Sea Res., 47, 789–824,
10.1016/S0967-0637(99)00112-0, 2000.
van Aken, H. M. and Becker, G.: Hydrography and through-fow in the north-eastern
North Atlantic Ocean: the NANSEN project, Prog. Oceanogr., 38, 297–346, 1996.von Schuckmann, K., Le Traon, P.-T., Alvarez-Fanjul, E., Axell, L., Balmaseda,
M., Breivik, L.-A., Brewin, R. J. W., Bricaud, C., Drevillon, M., Drillet, Y.,
Dubois, C., Embury, O., Etienne, H., Sotillo, M. G., Garric, G., Gasparin, F.,
Gutknecht, E., Guinehut, S., Hernandez, F., Juza, M., Karlson, B., Korres, G.,
Legeais, J.-F., Levier, B., Lien, V. S., Morrow, R., Notarstefano, G., Parent,
L., Pascual, Á., Pérez-Gómez, B., Perruche, C., Pinardi, N., Pisano,
A., Poulain, P.-M., Pujol, I. M., Raj, R. P., Raudsepp, U., Roquet, H.,
Samuelsen, A., Sathyendranath, S., She, J., Simoncelli, S., Solidoro, C., Tinker,
J., Tintoré, J., Viktorsson, L., Ablain, M., Almroth-Rosell, E., Bonaduce,
A., Clementi, E., Cossarini, G., Dagneaux, Q., Desportes, C., Dye, S., Fratianni,
C., Good, S., Greiner, E., Gourrion, J., Hamon, M., Holt, J., Hyder, P., Kennedy,
J., Manzano-Muñoz, F., Melet, A., Meyssignac, B., Mulet, S.,
Buongiorno Nardelli, B., O'Dea, E., Olason, E., Paulmier, A., Pérez-González,
I., Reid, R., Racault, M.-F., Raitsos, D. E., Ramos, A., Sykes, P., Szekely, T.,
and Verbrugge, N.: The Copernicus Marine Environment Monitoring Service Ocean
State Report, J. Operat. Oceanogr., 9, s235–s320, 10.1080/1755876X.2016.1273446, 2016.von Schuckmann, K., Le Traon, P.-Y., Smith, N., Pascual, A., Braseur, P., Fennel,
K., and Djavidnia, S.: Copernicus Marine Service Ocean State Report, Issue 2,
J. Operat. Oceanogr., 11, s1–s142, 10.1080/1755876X.2018.1489208, 2018.Zenk, W. and Armi, L.: The complex spreading pattern of Mediterranean Water off
the Portuguese continental slope, 37, 1805–1823, 10.1016/0198-0149(90)90079-B, 1990.
Zenk, W., Schultz Tokos, K., and Boebel, O.: New observations of meddy movement
south of the Tejo Plateau, Geophys. Res. Lett., 19, 2389–2392, 1992.