OSOcean ScienceOSOcean Sci.1812-0792Copernicus PublicationsGöttingen, Germany10.5194/os-14-971-2018A model perspective on the dynamics of the shadow zone of the eastern tropical North Atlantic – Part 1:
the poleward slope currents along West AfricaModeling the West African boundary current – Part 1KountaLalasoxnalala@gmail.comCapetXavierJouannoJulienKolodziejczykNicolashttps://orcid.org/0000-0002-0751-1351SowBamolGayeAmadou Thiernohttps://orcid.org/0000-0002-3688-1351Laboratoire de Physique de l'Atmosphère et de l'Océan Siméon Fongang, ESP
/UCAD, Dakar, SenegalLOCEAN Laboratory, CNRS-IRD-Sorbonne Universités-MNHN, Paris, FranceLEGOS Laboratory, IRD-Univ. Paul Sabatier-Observatoire Midi-Pyrénées, Toulouse, FranceLaboratoire d'Océanographie Physique et Spatial, IFREMER-IRD-CNRS-UBO, IUEM, Plouzané, FranceLaboratoire d'Océanographie, des Sciences de l'Environnement et du Climat, UASZ, Ziguinchor, SenegalLala Kounta (soxnalala@gmail.com)10September201814597199712February201819March201812July201823July2018This work is licensed under the Creative Commons Attribution 4.0 International License. To view a copy of this licence, visit https://creativecommons.org/licenses/by/4.0/This article is available from https://os.copernicus.org/articles/14/971/2018/os-14-971-2018.htmlThe full text article is available as a PDF file from https://os.copernicus.org/articles/14/971/2018/os-14-971-2018.pdf
The West African seaboard is one of the upwelling sectors that has received
the least attention, and in situ observations relevant to its dynamics are
particularly scarce. The current system in this sector is not well known and
understood, e.g., in terms of seasonal variability, across-shore structure,
and forcing processes. This knowledge gap is addressed in two studies that
analyze the mean seasonal cycle of an eddy-permitting numerical simulation of
the tropical Atlantic. Part 1 is concerned with the circulation over the West
African continental slope at the southernmost reach of the Canary Current
system, between ∼8 and 20∘ N. The focus is on the depth range
most directly implicated in the wind-driven circulation (offshore and coastal
upwellings and Sverdrup transport) located above the potential density
σt=26.7 kg m-3 in the model (approx. above 250 m of depth). In
this sector and for this depth range, the flow is predominantly poleward as a
direct consequence of positive wind stress curl forcing, but the degree to
which the magnitude of the upper ocean poleward transport reflects Sverdrup
theory varies with latitude. The model poleward flow also exhibits a marked
semiannual cycle with transport maxima in spring and fall. Dynamical
rationalizations of these characteristics are offered in terms of wind
forcing of coastal trapped waves and Rossby wave dynamics. Remote forcing by
seasonal fluctuations of coastal winds in the Gulf of Guinea plays an
instrumental role in the fall intensification of the poleward flow. The
spring intensification appears to be related to wind fluctuations taking
place at shorter distances north of the Gulf of Guinea entrance and also
locally. Rossby wave activity accompanying the semiannual fluctuations of
the poleward flow in the coastal waveguide varies greatly with latitude,
which in turn exerts a major influence on the vertical structure of the
poleward flow. Although the realism of the model West African boundary
currents is difficult to determine precisely, the present in-depth
investigation provides a renewed framework for future observational programs
in the region.
Seasonal climatology of Sverdrup transport (m2 s-1;
excluding the Ekman flow in the surface layer) computed from the DFS5.2 wind
forcing fields averaged for January–March (a), April–June
(b), July–September (c), and October–December
(a). The main regional flow, thermohaline features, and capes are
indicated in panel (a): the Canary Current (CC), North Equatorial
Current (NEC), North Equatorial Counter Current (NECC), north Equatorial
Undercurrent (NEUC), Guinea dome (GD), and Cape Verde frontal zone (CVFZ), which
separates the realm of the North Atlantic Central Waters (NACWs) and South
Atlantic Central Waters (SACWs). The TROP025 100 m isobath along which several
analyses are made is shown in black. Three geographical limits used to define
the integrated upwelling indices computed in Sect. are
shown with blue dots.
Introduction
The meridional extent of the Canary Current system (CCS) is one of its
remarkable features. The northern (southern) extreme is the northern
tip of the Iberian peninsula at ∼40∘ N (Cape Roxo at
∼12∘ N). Between approximately 35 and
20∘ N the system is aptly named. The Canary Current is the slow
southward return flow of the North Atlantic subtropical gyre flowing offshore
of northwest Africa (see setting in Fig. a). About
Cap Blanc (21∘ N) the Canary Current bifurcates toward the
southwest, away from the African continent, and feeds the North Equatorial
Current (NEC). Between ∼20 and 12∘ N, the southern end of the
CCS is well separated from the northern CCS (nCCS) by the Cape Verde frontal
zone (CVFZ) along which flows the NEC. Down to ∼200–300 m, major
contrasts exist across the CVFZ in terms of water masses: North Atlantic
Central Water (NACW) to the north and fresher South Atlantic Central Water
(SACW) to the south. At depths greater than ∼300 m, NACWs are found
further south and the water mass contrast fades away
e.g.,. Note that water masses are
traditionally separated into surface waters (potential density anomaly
σt lower than 26.3), upper central waters (26.3<σt<26.8),
and lower central waters (26.8<σt<27.15)
. South of the CVFZ,
the southern Canary Current system will in this study be referred to as
the eastern tropical North Atlantic (ETNA) to underscore its distinct character
and avoid overemphasizing oceanic connection and dynamical similarities with
the nCCS (to the contrary we will highlight the importance of the connections
with the tropical sector situated further south).
As we define it the ETNA is further delimited by the West African (WA) shores
and to the south by the Northern Equatorial Counter Current (NECC), which is
surface intensified and feeds the area with waters of equatorial origin
. The latitudinal position of the NECC
undergoes seasonal fluctuations as a consequence of the shift in the
Intertropical Convergence Zone (ITCZ) and trade wind position
. The NECC northernmost position at
∼10∘ N is reached in late summer–early fall as the flow reaches
maximal intensity. The wind regime exhibits important contrasting specificities
in the nCCS and ETNA. From the vicinity of Cap Blanc up to
∼25–30∘ N, wind is upwelling favorable all year. Further
north, upwelling winds are increasingly restricted to the summer period.
Conversely, the upwelling season is limited to the winter–spring period
between November and May in the ETNA, albeit less so when approaching Cap
Blanc. Important contrasts are also found in terms of wind stress curl (WSC;
not shown, but Sverdrup transport presented in Fig.
largely reflects WSC). Except nearshore where coastal wind drop-off can be
responsible for positive values, WSC is robustly negative over the nCCS
, which belongs to the North Atlantic subtropical gyre.
Conversely, ETNA WSC is predominantly positive because sea–land contrasts and
the shape of the African continent produce a curvature of the trade winds
favorable to cyclonic rotation, and also because the ETNA is a transition
region toward the ITCZ (i.e., the trade wind intensity gradually drops
southward).
East of 23∘ W, the ETNA has historically received limited attention
compared to the northern part of the CCS (and other eastern boundary regions)
and the regional circulation still suffers from important knowledge gaps.
Recently, the issue of the maintenance and possible expansion of the North
Atlantic deep oxygen minimum zone has prompted some studies concerned with
the density range 26.5–27.2 within
which low dissolved oxygen concentrations are found. In this study our focus
will be on the ETNA circulation and dynamics in a distinct, slightly lighter
density class σt≤26.7. This density class, straddling the
so-called surface and upper central water ranges, is important because it
feeds the coastal upwellings present along Senegal, Gambia, Mauritania, and
the southern part of Morocco . As part of a research
effort aimed at implementing an ecosystem approach to managing the WA marine
environment and fisheries, we are concerned with the origin of those upwelled
waters, the pathways they follow to reach the WA shore, and the dynamics that
underlie the existence of these pathways. In addition, note that the
relatively low dissolved oxygen concentrations found in this density range
have important implications since they contribute to anoxia or hypoxia over the
WA continental shelves through coastal upwelling .
This element of the biogeochemical context is an important motivation for this
work.
Model monthly climatology of zonal velocity at the ocean surface (m s-1).
This figure should be compared with Fig. 6 in .
Meridional (zonal) section at 26∘ W (a, b) (13∘ N; c, d) for temperature
(∘C) averaged during September–October. Observations and models respectively correspond to panels (a, c) and (b, d).
The ETNA broadly coincides with the shadow zone of the North Atlantic
subtropical gyre. In classical wind-driven circulation theories it is a place
of weak circulation, owing to the no-flow condition at the eastern
boundary . This means that thermocline waters, including
those in our density class of interest, are not directly ventilated even
though they outcrop overwhelmingly north of 20∘ N in the negative
WSC region . Strictly speaking though, the ETNA is not
part of the subtropical gyre. As mentioned above, it is characterized by
regional positive WSC, positive Ekman pumping, and, in virtue of the Sverdrup
relation, northward flow (the vertical distribution of which is not well
known and will be an important aspect of the present work). Two main
dynamical features have been identified in the region that are consistent
with these expectations: the Guinea thermal dome and the Mauritanian current.
The Guinea dome has been described in numerous studies
as a permanent quasi-stationary feature on the
eastern side of a quasi-zonal thermal ridge present over most of the basin at
∼12–14∘ N. The dome is characterized by a rise of isotherms in
the depth range 50–300 m. relate the Guinea dome to
the cyclonic rotation of the NECC when approaching the eastern end of the
basin, from eastward to northward and then westward as the flow connects to
the NEC. Their claim is that the quasi-zonal thermal ridge associated with
the NECC is reinforced by this cyclonic rotation, thereby giving rise to the
dome structure. Note, though, that the thermal ridge is much more visible (in
meridional cross sections; Fig. a, b) than the thermal
dome (in zonal cross sections; Fig. c, d). Despite its
supposed importance in conveying waters rich in dissolved oxygen toward the
North Atlantic OMZ , the Guinea dome remains to this date an
elusive circulation feature, with limited and contradictory results on its
position, dynamics, and variability, as this overview of the literature
indicates. analyze the Guinea dome structure and seasonal
variability using in situ observations and a primitive equation model. Based
on temperature distribution and geopotential anomaly fields, they conclude
that the Guinea dome is a permanent feature with some seasonal variability,
the upper thermocline center of the dome being found at about 9∘ N,
25∘ W in summer and 10.5∘ N, 22∘ W in winter. Their conclusions
partly contradict earlier studies that could not identify the upper
thermocline expression of the dome in winter and the
issue has not been settled since then . Finally, note that
ADCP measurements or averaged float
drifts show weak signs of westward flow on the northern
side of the Guinea dome in contrast to many schematic representations of the
circulation associated with the Guinea dome
.
Modeling has not led to much clarification, perhaps because the Guinea dome
is seldom reproduced with fidelity. The doming of the isopycnals along zonal
sections is clearly insufficient in the models of ,
, , and in ours (see below). In
contrast, OFES simulations presented by tend to overestimate
the doming (see their Fig. 2c and d). More problematically, these simulations
exhibit ETNA thermohaline uplifts that have distinct characteristics compared
to the real Guinea dome. For instance, in the subsurface
temperature field exhibits a cold coastal tongue between 10 and
25∘ N that protrudes offshore around 15∘ N (their Fig. 3). The cold tongue is quasi-stationary, while the protrusion is subjected to
an important seasonal modulation. Neither the model cold tongue nor the
protrusion can unambiguously be identified to match the observed Guinea dome
structure.
In terms of dynamical interpretation, the Guinea dome is being systematically
related to WSC forcing, but two variant explanations can be found in the
literature: local forcing or large-scale
forcing with some Rossby wave effects . Historically, this
disagreement has flourished in a context of large uncertainties on the WSC
patterns . The QuikSCAT climatology now allows us to
unambiguously demonstrate that the Guinea dome position does not coincide
with that of a WSC extremum so the dynamical rationalization for the
existence and position of the dome remains to be improved. In Part 2 the
Guinea dome will be considered in the regional circulation context. Herein we
focus on the flow over the WA continental slope (also referred to as coastal
flow given the regional perspective of this work) in the ETNA region, i.e., between approximately 8 and 20∘ N.
In this latitude range, the existence of upper ocean poleward currents over
the WA continental slope has been reported for a long time but only the basic
aspects of their structure (vertical and horizontal) and seasonal variability
are known. Two seasons of intensified poleward flow can be inferred from the
literature . During the upwelling season
(winter–spring), a poleward undercurrent naturally develops as also found in
the other upwelling systems . In summer–fall
another poleward flow intensification occurs . In
contrast to the earlier one, the flow is surface intensified and the surface
part of the flow is sometimes referred to as the “Mauritanian current”
following . This characteristic and an approximate
coincidence in time has led to the suggestion that this poleward flow pulse results
from the bifurcation of the summertime northern branch of the NECC as it
approaches WA . But the
specifics of the flow bifurcation (from zonal to meridional) have, to our
knowledge, never been described in dynamical terms. Alternatively or
complementarily, some authors invoke upwelling wind relaxation south of
∼21∘ N as the driving process for the
summer–fall pulse of poleward flow. Older studies tend to
insist on an origin in the Gulf of Guinea for the summer pulse
but this process has not been revisited
for a long time. In his 1989 review study Barton qualifies the knowledge of
the poleward undercurrent along the eastern boundary of the North Atlantic as
sketchy: “the arguments for its existence as a continuous entity are based
upon relatively few direct current observations, some interpretations of
temperature and salinity data, and a degree of speculation”. Almost 30 years
later, the situation is virtually unchanged with only a few irregular ship
ADCP transects to describe the boundary current system
. In particular, the works of
, , , and a few
follow-up and review studies remain the main sources of
observational knowledge about the WA slope currents between 10 and
20∘ N (between 5 and 10∘ N observations are even fewer).
In this context the point raised by about whether the
poleward undercurrent and the more seasonally intermittent surface
countercurrent (also called the Mauritanian current) are dynamically distinct
entities is still pending. The use of two different names to describe the
poleward flow depending on depth range implicitly suggests they are
dynamically distinct. We instead will prefer to use the unique and neutral
terminology “WA poleward boundary current” (or WABC in short) to refer to
the northward flow present over or in the vicinity of the WA continental
slope.
This overview strongly suggests that clarifications are needed on ETNA
circulation and dynamics. The present study and a companion paper (Part 2)
are a contribution to this needed effort. The focus is on waters within the
density class that most strongly responds to wind forcing because they are
transported by the Sverdrup flow and/or actively contribute to upwelling in
the ETNA sector, driven by Ekman pumping or coastal divergence. Due to the
sparseness of observations in this region modeling can be an invaluable
source of information. Our approach will essentially be based on the careful
analysis of an eddy-permitting NEMO model simulation see
Sect. . In this Part 1, the realism of the modeled
circulation and thermohaline structure is evaluated in Sect. 3 and will be
deemed sufficient to inform several related aspects of the WA ocean
dynamics. The seasonal cycle of the WABC will then be presented
(Sect. ). Its underlying dynamics will subsequently be explored
and discussed over the continental slope in relation to wind-forced coastal
trapped wave theory (Sect. ) and offshore in relation to
Rossby wave theory (Sect. ). In light of these
results and interpretations, a general assessment of the knowledge, knowledge
gaps, and model biases pertaining to the WA boundary current will be proposed
(Sect. 7). The source pathways for WA coastal upwelling waters and the broader
regional circulation (which turns out to be of key relevance to understanding
these pathways) will be examined in Part 2.
Data and methods
In this study, we use a numerical model, the oceanic component of the Nucleus
for European Modelling of the Ocean program NEMO3.6;. It
solves the three-dimensional primitive equations discretized on an Arakawa
C grid at fixed vertical levels (z coordinate). The grid horizontal resolution
is 1/4∘ and the configuration (referred to as TROP025 hereafter)
generously covers the tropical Atlantic
(35∘ S–35∘ N, 100∘ W–15∘ E). TROP025 has 75 vertical
levels, 12 (24) being concentrated in the upper 20 m (100 m). It
is forced at its lateral boundaries with daily outputs from the MERCATOR
global reanalysis GLORYS2V3 . The open boundary conditions
radiate perturbations out of the domain and relax the model variables to 1-day averages of the global experiment. Details on the numerical methods are
given in . At the surface, the atmospheric fluxes of
momentum, heat, and freshwater are provided by bulk formulae
. The simulation is forced with the Drakkar Forcing Set
DFS5.2 , which is mainly based on the ERA-Interim reanalysis
. DFS5.2 consists of 3-hourly fields for wind speed,
atmospheric temperature, and humidity and daily fields for longwave radiation, shortwave radiation,
and precipitation.
Some model–data comparisons are available in and
. Additional evaluation directly related to this study is
presented in the next section. It largely relies on the gridded version of
the Coriolis dataset for reanalysis version 4.2 (hereafter CORA4.2,
http://www.seanoe.org/data/00351/46219/, last access: 18 August 2018) for potential temperature and
salinity over the period 1990–2014. Its resolution is 0.5∘×0.5∘ but
the gridding software ISAS involves correlation length
scales that far exceed this mesh grid size, which is detrimental to the
representation of boundary currents such as the one we are interested in.
CORA4.2 includes the vast majority of available ARGO profiles and offers a
state-of-the-art description of the recent oceanic thermohaline state. The
model solution corresponds to a longer period (1979–2015) but this is deemed
inconsequential given the limited regional low-frequency variability. In
addition, note that a 15–20 % fraction of the CORA bins within 1000 km from
the WA shore have their monthly climatology built with less than 20 T–S vertical
profiles (not shown). There are thus substantial uncertainties in the true
ETNA climatological state irrespective of the period. For more information on
CORA4.2 readers are referred to .
Mathematical symbols have their usual meaning. T, S, σt, and ρ
respectively refer to potential temperature, salinity, potential density
anomaly, and in situ density. The variables x (y) and u (v) refer to zonal
(meridional) directions and velocity. At any (x, y) position, we define
the depth anomaly δzT0 for a specific isotherm T0 as the depth
of that isotherm minus that of its long-term climatological mean.
Many diagnostics involve vertical integration between the surface and the
isopycnal surface σt=26.7 kg m-3. U26.7 and V26.7
are vertically integrated zonal and meridional transports over that depth
range. The geostrophic part of the transport is noted with a “g” subscript.
At this stage the choice of 26.7 may seem arbitrary but it will be justified
by several model analyses below and in Part 2. In particular, it will be
shown that the layer above σt= 26.7 includes an overwhelming
fraction of the Sverdrup and upwelling circulation in the model. σt=26.7 is also convenient because subsurface waters lighter than this value are
overwhelmingly of the SACW type, in contrast to deeper waters
.
For reference, the geostrophic part of the Sverdrup volume transport
denoted Vsv is defined as Vsv=fβ∂xτyρ0f-∂yτxρ0f,
where f is the Coriolis parameter, β=∂yf its derivative with
respect to the meridional coordinate, ρ0 is a reference density equal
to 1025 kg m-3, and (τx, τy) is the surface wind stress (in N m-2).
Dynamic heights are presented in Sect. . They are calculated at
50 dbar relative to 500 dbar according
to the standard formula ΔD50/500=g∫50050(ρ(T,S,p)-1-ρ(0,35,p)-1)ρ0dp,
with the additional assumption that pressure and depth are equivalent. The
associated geostrophic flow is classically given by
ug=gρ0f∫50050∂ρ∂ydz,vg=-gρ0f∫50050∂ρ∂xdz.
Potential vorticity (PV) is examined in the context of Rossby wave dynamics
in Sect. . It is expressed in density coordinates for a
fluid layer of thickness h between the potential density surfaces
σ1 and σ2:
PVσ1σ2=f+ξh≈fh,
where ξ is the vertical component of relative vorticity for the flow in
this fluid layer, which can be neglected given the smallness of the Rossby
number associated with the eastern boundary conditions under consideration.
In several instances, we wish to rotate the flow or wind at the model shelf
break to isolate their alongshore and/or along-slope component. To do so velocities
are rotated with respect to the orientation of the shelf break. This
orientation is computed at every grid point following the 100 m isobath
using centered differences. A total of 15 passes of a three-point filter with coefficients
(0.25, 0.5, 0.25) are subsequently applied to ensure some smoothness to the
angle used for the along-slope projection.
Model evaluation
In this section we carefully evaluate our simulation with respect to
available observations, mainly from the CORA4.2 database. As mentioned in the
previous section, time periods for the model and observations do not match
precisely. In addition observations are not particularly dense in our region
of interest (although many of the bins have ∼100 measurements per
climatological month) so we are concerned with a qualitative assessment of
model realism.
We start by comparing the monthly climatology of surface zonal currents in
TROP025 with the climatology derived from ARGO drifts obtained by
. The agreement is quite remarkable both in terms of
spatiotemporal patterns and current magnitude (compare Fig.
with their Fig. 6). The model captures the northern and equatorial branch of
the South Equatorial Current whose separation is most clear in boreal spring,
as also found in the observations. More importantly for our study, the NECC
seasonal cycle is realistic both in terms of north–south displacement
(northernmost extension and widest latitude range in August–September) and
change in flow magnitude (swiftest currents >0.3 m s-1
found in July–August). Note though that peak NECC currents in TROP025 seem a
bit weaker than observed. The eastward-flowing Guinea current, whose seasonal
variability will turn out to be of relevance for the WABC, is also adequately
represented. It is intensified between 10 and 0∘ W with
a slight peak in boreal summer and a marked decrease in flow speed from
September (when the Guinea current is strongest) to November. North of
10∘ N within 5–10∘ from the West African coast observed
zonal velocities are weak and variable but generally oriented westward. This
westward tendency is less marked for model velocities. This discrepancy may
at least partly be related to Stokes drift, which is mainly zonal, can reach
about 0.05 m s-1 in that region, and affects surfacing ARGO floats
.
Outside the equatorial region where TROP025 behaves adequately
we cannot more precisely evaluate the model circulation
from direct observations but the equatorial and tropical Atlantic thermohaline
structure is overall quite well represented. This is evident from a
model–data comparison of climatological temperature along two vertical
sections (Fig. ) at 13∘ N and 26∘ W. The
latter crosses the thermal ridge associated with the NECC. The position of
the zonal section is chosen at the latitude typically associated with the
Guinea dome. Following and these fields are
shown for the September–October period during which the Guinea dome is supposed to
be most marked, but model–data agreement and differences are quite similar
for other months. Along 26∘ W thermocline displacements with latitude are
quite faithfully reproduced. For example, the deepening of the 20 ∘C
isotherm from 12∘ N at the top of the thermal ridge to
20∘ N is ∼55 m in the observations vs. 50 m in TROP025.
The main bias concerns the sharpness of the thermocline, which is
insufficient, presumably as a consequence of an overly strong model diapycnal
diffusivity. As a consequence, temperature is 1.5 ∘C too high at 100 m
of depth in the vicinity of the thermal ridge. A similar bias is found along
13∘ N where the model thermocline is too diffuse. Along that
section, the dome structure is manifest in the model with a deepening of the
isotherms colder than ∼20∘C toward the coast, albeit with less
amplitude than found in the observations. As in the observations the
longitude east of which the deepening occurs shifts westward for colder
isotherms. On the other hand, TROP025 is unable to produce the limited
vertical displacement observed east of 20∘ W for isotherms above 20∘. So
overall, the modeled Guinea dome present in TROP025 is weaker than observed.
This model bias is common , with the notable
exception of . We suspect that this bias is also reflected in the
intensity of poleward velocities that may be underestimated in TROP025.
Summer (June–September, a, b) and winter (December–February,
c, d) mean dynamic height at 50 m relative to 500 m (m2 s-2). The
contour interval is 1 m2 s-2. The associated geostrophic
circulation is also shown (vectors). Panels (a, c) are
for observations (TROP025). To ease comparison with the observations,
3 m2 s-2 was uniformly subtracted from the model
fields.
Depth (m) of the isopycnal density surface σt=25.2(a, b) and
σt=26.7(c, d) in CORA (a, c) and TROP025 (b, d).
Seasonal climatologies of dynamic height ΔD50/500 (see
mathematical definition in Sect. ) shown in Fig.
for winter and summer confirm this suspicion. Model and observation general
patterns are consistent with each other on the NEC and NECC signatures and
their winter–summer changes. The position and intensity of the NECC is quite
similar over most of the domain for both seasons, particularly in summer. In
winter, the precise form of the cyclonic circulation around the Guinea dome
is less well reproduced by the model. The main differences are east of
24∘ W where the model NECC exhibits some meanders not seen in the
observations. A northward branch is evident in the model and observations but
the model flow turns more gradually, starting from further to the west at
∼25∘ W than in the observations (e.g., compare geopotential
lines 63 in summer). East of 20∘ W the tilt of the geopotential
lines is noticeably different in the model and the observations. The regional-scale meridional gradient of geopotential in the ETNA is significantly
stronger in the observations than in TROP025, respectively 7 vs. 4 m2 s-2 in summer and 6 vs. 2.5 m2 s-2 in winter. Implications of
this bias will be discussed in Part 2, but we note again that model regional
circulation is in good qualitative agreement with observations.
Our analyses will systematically use the isopycnal surfaces σt=26.7
and, more infrequently, σt=25.2 because they approximately limit
the range of subsurface waters involved in the model WA coastal upwellings
and meridional transport. In Fig. we show the climatological
depth of these two isopycnal surfaces in the model and CORA observations.
Qualitative agreement between the two is evident, e.g., in terms of
outcrop line position, east to west deepening tendency for σt=25.2,
and shape and amplitude of the σt=26.7 doming in the central part of
the basin. Very close to shore along WA the model σt=26.7 isopycnal
surface is not close enough to the surface, presumably because our
eddy-permitting resolution is insufficient to adequately resolve the coastal
upwelling per se . However, this bias is limited to
∼20 m and is not crucial for the present continental slope and open ocean
investigation.
In attempting to explain the WABC seasonal cycle we will focus on the
dynamics along the coast of the Gulf of Guinea. Therefore, a model evaluation is
conducted at 4∘ N, 5∘ W, i.e., over the continental slope
south of Abidjan, Ivory Coast. The seasonal cycle of temperature T(z) at
this location was reported by for the period 1957–1964 (see
his Fig. 15b). Comparison with our Fig. reveals the good
level of realism of TROP025. Most noticeably, the model produces a
semiannual oscillation of the thermocline that is most pronounced in the
depth range 50–150 m with the highest (lowest) temperatures reached in
April and October–November (August and January). Model oscillations resemble the
observations in terms of phase, upward propagation of the summer–fall doming
tendency, and contrast in amplitude between the winter–spring (weak) and
summer–fall (strong) oscillations. On the other hand, a quantitative
difference concerns the amplitude of the oscillations, which are
underestimated by 50 % or more in the model, e.g., 30 m (50 m)
peak to peak amplitude for the seasonal cycle of δz16 in TROP025
(in the observations). This bias is particularly pronounced below 150 m. Finally, we note that the marked deepening phase between July and October is
better captured than the preceding shallowing phase. The latter occurs more
rapidly in the observations, e.g., over a 2-month period in
June–July for the 16 ∘C isotherm vs. 3–4 months in the model. To put
these biases into perspective it would be useful to know the degree to which
the very local conditions at 4∘ N, 5∘ W that TROP025 does not
represent (e.g., fine-scale irregularities of the shoreline or
continental shelf–slope bathymetry) contribute to the observed seasonal
cycle.
Time–depth representation of TROP025 climatological temperature
(∘C) at 4∘ N–5∘ W over the period 1982–2012. This
figure should be compared with the climatological cycle observed at the same
location between 1957 and 1964 his Fig. 15b. To evaluate
uncertainty 24 model climatologies were computed using 8-year
averaging periods as in the observations (with starting years from 1982
to 2005); the shallowest and deepest positions of four isotherms (11, 15, 18, and
21 ∘C) over this ensemble are shown with gray shading and indicate
the limited role played by interannual variability, particularly for the 18
and 21 ∘C isotherms.
Overall, the model climatological traits and dominant patterns of seasonal
variability are quite realistic, both at basin scale and more locally in the
ETNA. Our conclusion is thus that the model circulation and thermohaline
structure possesses a sufficient degree of realism to warrant further
in-depth analysis. In our discussion of the real ETNA dynamics and
circulation (Sect. and likewise in Part 2) we will keep
in mind the biases that have also been identified, including the relative
weakness of the model poleward flow along WA.
Seasonally averaged vertical–zonal section of meridional
velocity (in colors; cm s-1), σt (blue lines; kg m-3),
and mixed layer depth (black line; in meters) averaged between 13 and
15∘ N.
(a, c) Time–longitude diagram of vertically integrated
meridional transport (m2 s-1) averaged along-slope between
13 and 15∘ N (a) and between 7 and
9∘ N (c). Propagation speed at 3.5 cm s-1(a) as well as 7.4 and
3 cm s-1 (c) is shown with dashed lines.
These values are the ones we choose as most appropriate to describe the
propagation of patterns in the diagrams. (b, d) Associated climatologies of
meridional transport integrated vertically and across-shore. Vertical
integration is performed from the isopycnal surface σθ=26.7
up to the bottom of the mixed layer (a, c and red curve in b, d)
or up to the surface but excluding Ekman transport; i.e., we only take into
account the geostrophic flow (blue curve in b, d). Across-shore
integration is performed from the shoreline to the first location where
vertically integrated transport vanishes with a maximum longitude range of
3∘ (integration limit is indicated with the thin solid line in a and c).
Seasonally averaged vertical section of along-slope velocity (in colors; cm s-1), σt (blue lines; kg m-3), and mixed layer
depth (black line; in meters) following the shelf break (across-shore averaging between the 100 m isobath and
six grid points 150 km offshore).
Monthly climatology of meridional transport (m2 s-1)
integrated between the isopycnal surfaces σt=26.7 and the
surface, excluding wind-driven Ekman transport calculated from TROP025 wind
fields. The two thin dashed lines represent the location where zero PV
gradients are found in Fig. .
The seasonal cycle of the West African boundary current
Poleward boundary currents are ubiquitous along eastern boundary continental
slopes , particularly those subjected
to upwelling-favorable winds for which subsurface undercurrents are essential
flow features . WA is no
exception . Poleward currents are present in TROP025 as
revealed by the seasonally averaged zonal transects of v (Fig. ).
Figures – offer complementary views of the
structure and seasonality of the WABC.
At 14∘ N (i.e., at a central location in the ETNA) a poleward
undercurrent is visible over the continental slope for all seasons except in
summer (July–September) but it is most marked in fall and to a lesser extent in
spring (Fig. ; hereafter we refer to these two poleward flow
intensification periods as Pf and Ps, respectively). The undercurrent
appears to be strongly baroclinic with deviations of the isopycnals changing
sign in the vertical: upward toward the shore above ≈75 m of depth
and downward below. Isopycnal displacements reach ∼100 m for the 26.7
isopycnal between 25 and 17.5∘ W. The core of the
undercurrent is located at 50 to 100 m of depth with peak velocities reaching
6–8 cm s-1. In fall surface currents are oriented poleward. The absolute
flow maximum is found at ∼18∘ W and coincides with the mixed
layer base. This surface-intensified flow is the model “Mauritanian
current”. A near-surface secondary maximum is also present in spring at
approximately the same longitude and depth. In winter and summer a core of
weak poleward flow present a few hundred kilometers from shore is suggestive
of the radiation of westward-propagating Rossby waves from the continental
slope as also found in other regions
, particularly in the tropics. This
is confirmed with a time–longitude diagram of vertically integrated
meridional geostrophic transport Vg26.7 at about 14∘ N
(integration bounds for the integral follow the isopycnal surface σt=26.7 and surface; see Sect. ). The former broadly coincides
with the bottom of the poleward flow. The diagram (Fig. a)
exhibits clear signs of westward propagation with a speed around 3.5 cm s-1 and a dominant wavelength of about 650 km. The signal amplitude
decreases dramatically over 3–5∘ of longitude. Similar or even
shorter scales of attenuation are obtained for other semiannual Rossby
signals emanating from eastern boundary systems .
At 8∘ N (i.e., the southern end of our study region), the time–longitude
variability of the meridional transport is more complex. Two propagation
speeds can be identified (Fig. c) but the patterns are not
those expected from a simple superposition of two waves. The processes at
stake in the generation of the ETNA Rossby wave field and the possible
reasons underlying their rapid offshore attenuation at 14∘ N are
discussed in Sect. .
As a future point of comparison to other transport estimates, we compute the
meridional transport vertically and zonally integrated at 8 and
14∘ N. Zonal integration is performed from the coast to the first
offshore location where the flow changes direction with a maximum longitude
range of 3∘, so the width over which this transport is achieved varies in
time. At 14∘ N, the flow is poleward all year except for two
brief periods of weak equatorward flow in January and July. The poleward
transport along the WA boundary is seasonally variable with peak values
reaching 2 Sv or more during the two peak seasons in May–June and
September–November, with differences between Ps and Pf being around 20 %
(2 Sv in spring vs. 2.5 Sv in fall). Compared to 14∘ N, the
transport seasonal cycle at 8∘ N is more symmetric around 0
(equatorward flow is marked in summer and late fall to early winter; spring
intensification is weaker) with the notable exception that the fall
intensification reaches similar values in both latitude ranges.
Along-slope vertical sections of seasonally averaged along-slope current are
shown in Fig. . For each latitude, the current intensity is
obtained by across-slope averaging the along-slope flow between the 100 m
isobath and 150 km offshore (six grid points); i.e., Fig.
is representative of the flow over the continental slope. The regional-scale
coherence of the WABC is clearly visible although some minor flow
discontinuities result from meandering and eddy formation in the
vicinity of the major capes as better seen in Fig. see also. The northern bound of the poleward flow
varies significantly between Ps and Pf
(20∘ N vs. 25∘ N, respectively). The surface flow is
stronger during the latter but poleward currents are otherwise found over a
relatively similar depth range that deepens poleward, being located above
100 m (250 m) of depth south of 10∘ N (between
10 and 20∘ N). In fall when the poleward flow reaches
further north it extends down to 350–400 m north of 20∘ N. Note
that for latitudes between 10 and 20∘ N, σt=26.7
corresponds quite accurately with the bottom depth of the WABC, which partly
motivated our choice of this isopycnal.
In winter weak but coherent poleward flow is still present over the
latitudinal range 7–15∘ N. Equatorward flow is mainly found north
of 20–22∘ N in the nCCS (where the Canary Current hugs the coast)
and in the subsurface below the WABC during Ps and Pf. This highlights
the importance of baroclinic effects in the dynamics of the flow. In winter
and summer intense near-surface equatorward flow is found south of ≈10∘ N, but
it is confined to within 50 m of the surface.
The spatiotemporal complexity of the WABC behavior is further revealed in
Fig. , which shows vertically integrated geostrophic flow
from the surface down to σt=26.7. We make several important
observations. First, the general coherence of the flow in the meridional
direction (timing of the poleward flow intensifications and their relaxation)
is noticeable as is a general westward propagation tendency. The
meridional flow is organized in strips that appear at the coast and move
offshore. The strips become increasingly tilted in the northeast–southwest
direction as they move offshore, consistent with a faster westward
propagation speed closer to the Equator. This is particularly evident in
March and October when examining the two “phase lines”
We use this
terminology in anticipation of an interpretation of this meridional flow
signal in terms of Rossby wave dynamics. Nevertheless, a tendency for flow
strips to disaggregate is also noticeable, e.g., in July when the
strip of poleward flow associated with Ps is broken into several rounded
pieces. This underscores the complexity of the dynamics and the possible role
of parallel flow instabilities in destabilizing the WABC.
corresponding to
Ps and Pf, separated by a thin band of equatorward flow. Propagation
becomes increasingly ambiguous when approaching Cap Blanc at about
20∘ N.
Finally, although this description of the WABC seasonal cycle strongly
suggests the importance of its semiannual cycle
The relative
importance of the semiannual frequency was more precisely quantified via a
harmonic analysis performed at each grid point for the time series of monthly
averaged meridional velocities over the period 1982–2012. Over the entire
depth range above σt=26.7 the amplitude associated with the
semiannual period is at least 50 % larger than that for the annual period
in most of the ETNA and several-fold larger over the continental slope where
the WABC occurs.
, note that a perfect semiannual oscillation would
translate into an exact correspondence between the left and right panels in
Fig. . In contrast, the winter time interval from Pf to
Ps appears to be a bit shorter than the summer interval from Ps to
Pf (e.g., poleward transport is present nearshore in February but still
absent in August). This is confirmed at 14∘ N by inspection of
Fig. a. Asymmetry between the meridional flow during Ps and
Pf is more generally confirmed by Fig. b, which reveals a
sharper peak of northward transport for the latter period.
The WABC coastal dynamics
In the ETNA, positive WSC input is a priori a fundamental ingredient in the
generation of poleward flow both nearshore (Fig. ; see , for similar effects
in other eastern boundary systems)
and at larger scale
Fig. and Part 2. To be
more quantitative, we compare the theoretical Sverdrup transport Tsv and
geostrophic WABC transport in TROP025 over the continental slope
(Table ). Geostrophic model transports above 500 and 1000 m
of depth, as more commonly estimated in past studies
e.g.,, are also given. Results differ
strongly between the southern, central, and northern ETNA. At 14∘ N, all
model estimates correspond to over 75 % Tsv with limited changes when
increasing the range of integration. In contrast, model transport increases
steadily with the range of integration at 20∘ N, reaching 70 %
Tsv when integration goes down to 1000 m. At 8∘ N Tsv is
significantly weaker (see also Fig. ). There, the
model transport also increases steadily with depth but with values in
systematic excess of the Sverdrup transport estimate, reaching twice Tsv
when integration goes down to 1000 m.
Climatological Sverdrup transport Tsv (geostrophic part,
see Sect. ) computed from DFS5.2 winds over across-shore sectors
from the shelf break to 150 km offshore for three different latitudes ranges
(2∘ wide centered over the latitude reported in the first column).
Model geostrophic transports computed over the same areas are indicated as a
percentage of Tsv for three different ranges of vertical
integration from the surface to σt=26.7 (Vg26.7,
third column), 500 m (V500m, fourth column), or 1000 m (V1000m,
fifth column).
There are several reasons why the percentages in Table are not
strict determinations of the fraction of WABC transport that can be
attributed to WSC. First, the cross-shore width and transport of the WABC is
not uniquely defined because it varies as a function of latitude and time of
the year (see our estimation procedure used in
Fig. b). Bottom pressure torque can also cancel part of the WSC
contribution to the barotropic vorticity balance e.g.,in the context
of an eastern boundary current. In addition, momentum fluxes
by mesoscale eddies are known to redistribute WSC input, particularly in the
across-shore direction . Most importantly, the
Sverdrup balance is a constraint on the total barotropic flow. Thus, although
Sverdrup flow tends to be concentrated in the thermocline and above
, the WABC transport as we define it (above σt=26.7) does not solely reflect the Sverdrup balance, but also baroclinic
processes and how they vary in time (e.g., on seasonal scales;
Fig. ) and space (e.g., the meridional changes in
baroclinic structure; Fig. and Table ). In
this context and pending further progress with model sensitivity experiments
we hypothesize that mean poleward transport in the vicinity of the WA
continental slope arises from local wind stress curl driving Sverdrup flow
plus a combination of baroclinic response to remote wind forcing
as far as the equatorial band, baroclinic
response to meridional gradients of the Coriolis frequency
, and alongshore gradient in wind stress curl
.
We now turn to the seasonal cycle of the WABC about which more can be said
based on the TROP025 experiment alone. The main processes underlying
Ps-
and Pf-intensified poleward transport could a priori result from four
distinct (not mutually exclusive) processes: (i) local generation of a poleward
undercurrent in conjunction with variable coastal upwelling conditions, (ii)
remote forcing of poleward flow with subsequent propagation in the form of
coastal trapped waves, (iii) local modulation of the nearshore Sverdrup
transport in relation to the seasonal cycle of the wind stress curl, and (iv)
excitation of free Rossby wave modes at the semiannual frequency, e.g., by processes (i)–(iii) if their associated wavenumber–frequency match the
dispersion relation.
As mentioned above, large deviations from the Sverdrup balance are possible
at fine temporal scale (e.g., sub-annual) ,
particularly near lateral ocean boundaries . One striking
discrepancy between the WABC and corresponding Sverdrup transport concerns
their respective seasonal cycles that bear no resemblance, as illustrated in
Fig. at 14 and 20∘ N. Over the
continental slope at 14∘ N, Sverdrup transport is dominated by an
annual cycle that sharply contrasts with the semiannual cycle of
Vg26.7 (a similar contrast is found for the geostrophic
meridional transport computed with integration limits at 500 or 1000 m
of depth; not shown). At 20∘ N, the semiannual cycle of the meridional
transport is less prominent but differences with the local Sverdrup transport
remain important. These arguments lead us to exclude process (iii) as a
process responsible for the semiannual cycle of the WABC. In the remainder
of this section the respective roles of (i) and (ii) are considered, while the
possible role of (iv) will be discussed in the next section.
Seasonal cycle of Sverdrup transport
Vsv=fβcurl(τρ0f) (red lines) and
meridional geostrophic transport vertically integrated between σt=26.7 and the surface (blue lines). Both transports are across-shore
integrated between the 100 m isobath and six grid points offshore. Solid lines
are for the Cape Verde region (13–15∘ N), while dashed
lines are for the Cap Blanc region (19–21∘ N).
Time–along-slope distance diagram following the 100 m isobath (see
location in Fig. ) for the TROP025 climatological
seasonal cycle of (a) the depth anomaly of the 18∘ isotherm
δz18 (colors; in meters) and SLA (contours; cm), (b)
along-slope geostrophic transport between σt=26.7 and the surface
Vg26.7 (m2 s-1), and (c) across-shore Ekman
transport anomaly. Oblique solid black lines correspond to propagation speeds
of 0.2–0.3 (February–June) and 0.9 m s-1 (August–January). They are
subjectively drawn in (a) (and repeated in b and
c) to indicate where spatiotemporal patterns are suggestive of
along-slope propagation. In panel (c), vertical gray lines delineate
the sectors over which sector upwelling indices are computed (see
Fig. ). Power spectral densities associated with the
annual and semiannual harmonics of the wind stress curl and cross-shore
Ekman transport along the 100 m isobath are displayed above (a) and
(c), respectively.
The significance of semiannual fluctuations has been established for the
circulation of several regions of the equatorial and tropical ocean
. In a nutshell, this arises from having the ITCZ pass twice
a year over these regions. The central and eastern equatorial Atlantic
receives significant semiannual forcing from the winds
. Variability at this frequency is further enhanced by
abrupt temporal changes in zonal winds and basin-mode
quasi-resonance . The equatorial response can then
be propagated poleward along eastern boundaries by coastal trapped waves, as
occurs in the northern and southern Pacific , for
instance. The Northern Atlantic is peculiar in that coastal trapped waves
generated by the reflection of equatorial Kelvin waves in the eastern Gulf of
Guinea (GG) have to propagate along a long and corrugated stretch of
coastline to reach WA latitudes, with a significant part of the coastline
being situated on the edge of the equatorial band at ∼4∘ N.
This basin geometry is not prone to the coastal transmission of equatorial
signals . Despite some controversies, several
studies have dismissed a connection between the equatorial region and the
northern GG . To gain further insight into the
possible remote sources of semiannual poleward flow off WA, we computed
time–along-slope distance diagrams (i.e., following the continental
slope) for the climatological cycle of 18∘ isotherm depth anomalies
(δz18, a proxy for thermocline depth anomaly) and
Vg26.7 anomalies (Fig. ). The along-slope
coordinate covers from 4.2∘ N, 4.75∘ E (inside the GG) to
25∘ N, 16.5∘ W. A diagram for climatological anomalies of
across-shore Ekman transport is also presented.
Overall, we observe clear signs of long-range poleward propagation for
thermocline depth and along-slope flow but this assertion should be qualified
for the following reasons: (1) propagation is more clearly visible for δz18 than for
Vg26.7 (which is considerably noisier) and sea level anomaly
(SLA; contours in Fig. a) ; (2) propagation is more
evident during Pf than Ps, with distinct time–space patterns for the
two periods; (3) examination of Fig. reveals an
off-equatorial maximum in poleward flow and thermocline depression in the
longitude range 4–14∘ W (i.e., in the GG) with no clear
connection to the area east of 0∘ W; (4) the temporal lag between
thermocline depth anomaly (δz18) and geostrophic velocity
(Vg26.7) seasonal fluctuations is not consistent with Kelvin
wave theory; and (5) propagation becomes increasingly ambiguous north of Cape
Verde at ≈15∘ N.
To help with the discussion of reasons (2) and (3), Fig. displays
the seasonal cycle of the upwelling index (i.e., Ekman transport)
integrated over the stretch of coastline between 3∘ E and 7.5∘ W:
UIGG=1ρ0f∫0∘W10∘Wτxsinα-τycosαds,
where α is the angle between the north and the tangent to the
coastline leaving land on the right and s is the curvilinear coordinate
following the 100 m isobath. UIWAC (UIWA) is similarly
defined for the longitude (latitude) band 7.5–17∘ W (10–20∘ N), corresponding to the “West African corner”
between Cape Palmas and Cape Roxo (West Africa between Cape Roxo and
Cap Blanc; see Fig. for locations). All three
sectors are of comparable length.
Monthly mean climatology of across-shore Ekman transport (Sv)
integrated along the 100 m isobath for three sectors: in the Gulf of Guinea
between longitudes 3.5∘ E and 7.5∘ W (black; GG), further
west between 7.5 and 17∘ W (blue; WAC), and between 10 and
20∘ N (red; WA). The position of these sectors, which are of
comparable length, is indicated in Figs. a and
c.
Reason (1) is partly expected because along-slope velocities associated with coastal
trapped waves should be approximately geostrophic, and hence they depend on the
across-shore derivative of δz18. Note also that
standing meanders of the WABC past topographic irregularities can produce
substantial excursions of the flow away from the shelf break and hence rapid
along-slope changes of Vg26.7 (see Fig. ; the
position in the main capes is indicated in Fig. ). The
lack of clear propagation tendency found for SLA was previously noticed by
and reflects the fact that sea level over the slope area is
at least partly decoupled from δz18 (and also from
Vg26.7). This limits the utility of altimetry to investigate
remotely forced dynamics off WA.
Regarding reason (2), we note that south of 15∘ N a propagation phase speed
can be identified with reasonable confidence during Pf, especially for
δz18. We estimate c at ≈ 0.9 m s-1 in
Fig. a and this value is also applicable to
Vg26.7. This is compatible with low vertical-mode coastal
trapped wave propagation and, more importantly, consistent with the
propagation speed inferred by for the GG coastal upwelling
signals (0.7–0.8 m s-1).
Note that our value is ∼50 %
slower than the one found by in their numerical simulation
over a similar area. Possible reasons for this difference include numerical
differences in the grid resolution (higher in TROP025 by a factor 2 and 2.5
in the horizontal and vertical direction, respectively) and treatment of
viscosity .
On the other hand, the coherence of the signals expected from poleward
propagation is weaker around Ps, especially for Vg26.7. Even
for δz18 propagation speed is ambiguous and seems to change with
time (≈0.2 m s-1 at the transition between negative and
positive δz18 anomalies but a bit faster toward the downwelling
peak (∼0.3 m s-1), which roughly coincides with a sign change of
Vg26.7 north of Cape St Ann (Fig. ). Such
values are untypical for coastal trapped waves. The slow northward shift of
the downwelling signal (negative δz18 and strong poleward flow)
may alternatively be attributed to the progressive seasonal migration of the
upwelling wind region (related to the seasonal displacement of the ITCZ) but
the correspondence between panels (a), (b), and (c) in Fig. is only partially supportive of this. In addition,
propagation at ≈0.9 m s-1 may also be present, e.g.,
toward the end of Ps in May. This suggests that both local and remote
responses to winds combine to produce the winter–spring WABC intensification.
Examination of Fig. reveals a complex picture in which
each separate coastal sector contributes to Ps upwelling relaxation over a
slightly different time period: April to August for WA; March to June for
WAC; and March to May for GG. Also note that the GG relaxation is immediately
followed by marked increasing upwelling tendency from May to July. Overall,
forcings over the different sectors largely oppose each other; hence the
weakness of propagating oceanic signals and perhaps also the weakness of the
poleward flow relative to Pf.
With respect to Pf and reason (3), our analyses suggest the existence of a remote
origin for the WABC intensification off WA, with an evident off-equatorial
maximum in poleward flow and thermocline depression in the longitude range
4–14∘ W, i.e., in the GG (Fig. ). For the
period October–December, the largest positive values are found in this longitude range.
On the other hand, Hovmüller diagrams for δz18 and
Vg26.7 exhibit some pattern changes at ≈0∘ W
near the left edge of Fig. . We take this as an
indication that the equatorial region is not implicated in the generation of
the Pf CTW signal
The same is also true for the preceding
upwelling phase between June and September, in agreement with the conclusions of
and . In particular, δz18
minima reflecting summer upwelling tendency are much more pronounced between
4 and 10∘ W than near 0–2∘ E (Fig. a).
.
Examination of 2-D monthly regional maps for δz18 (not shown)
confirm the absence of oceanic signal propagation between the Equator and the
northern part of the GG. Figure confirms the importance
of the GG sector as a forcing region for poleward flow during Pf. An
abrupt upwelling relaxation takes place in the GG from August to November
when the ITCZ approaches and passes over this sector .
This relaxation is far steeper than the boreal winter one (compare the two
drops in upwelling index UIGG in Fig. and the
corresponding local thermal response in Fig. and
, his Fig. 15b). To our knowledge, there has been no
previous mention of the role played by the GG wind cycle as a source of
remote forcing for the poleward flow in the southern part of the Canary
Current system (although remote forcing from equatorial origin has been
invoked to explain the seasonal cycle of subsurface temperature off Dakar; ). UIWAC also decreases during
Pf,
so winds in the WAC sector must contribute to WABC intensification, but the
amplitude of the relaxation is smaller by a factor close to 4. The
relative importance of the remote forcing associated with each sector depends
on their along-slope decay scale, which is poorly constrained and may depend
on a number of factors. Limited insight into this question can be gained by
comparing Pf and Ps remote forcings.
During Ps the WAC wind relaxation is about twice as intense as during
Pf and combines (between April and June) with the local relaxation of WA
winds. However, the ocean response in terms of poleward flow is significantly
weaker than the one during Pf both in terms of current magnitude and
meridional extent. GG winds are thus plausibly instrumental in driving the
model WABC intensification in fall and, conversely, opposing intensification
during most of Ps. Further analyses will be needed to clarify this because
the seasonal cycle of other environmental parameters may also be involved,
e.g., the near-surface density gradient along the waveguide, which is
larger in spring than in fall (Fig. ).
With respect to reason (4), δz18 and Vg26.7 are not
precisely in phase as they are expected to be for theoretical Kelvin waves in
a model for a single baroclinic mode e.g., reduced gravity
model;. A phase shift of the order of 1 month exists
between the two variables with δz18 lagging; i.e., Pf
poleward flow intensification initiates while the thermocline is still in an
uplifted position (Fig. ). A similar discrepancy has been
noted before for the California undercurrent and attributed to the effects of
Rossby waves dynamics . Radiation of Rossby waves from the
coastal waveguide implies that pressure disturbances associated with CTW
dynamics propagate offshore. In turn, this modulates the along-slope flow,
which depends on the nearshore–offshore pressure difference. In particular,
this produces a phase lead for Vg26.7 compared to the
thermocline depth at the coast . We will discuss the Rossby
wave activity offshore of WA in the next section and add support to this
explanation.
With respect to reason (5), the propagation of thermocline depth anomalies associated
with Pf becomes progressively elusive beyond Cape Verde
Likewise,
lagged cross-correlation of the seasonal 18 ∘C depth between an origin
placed 4∘ E at 100 m of depth and all the other points along the 100 m
isobath degrades rapidly north of Cape Verde (not shown).
. This would not be
inconsistent with the major influence exerted by this cape on the poleward
flow and its dispersive effect on coastal trapped
waves . The area located between 15 and 20∘ N (i.e., Cape Verde and Cap Blanc, respectively) is also characterized by a
rapid shift in the dominant periodicity of δz18 and
Vg26.7 fluctuations (Figs. and
) from semiannual to annual. In particular, δz18 variability becomes increasingly complex with reduced magnitude when
approaching Cap Blanc where upwelling is permanent. Overall, TROP025
suggests the existence of a transition in this latitude range despite the
fact that the WABC can be present up to ∼25∘ N in fall.
Overall, no significant forcing in a semiannual period is present north of
Cape Palmas (see frequency decomposition of across-shore Ekman transport in
Fig. ) and our analyses indicate that the WABC seasonal
cycle is made of two parts that are distinct in terms of forcing mechanism.
The strongest poleward flow intensification occurs in fall, both in terms of flow
speed and also poleward extent (25∘ N vs. 20–22∘ N in
the model). Such differences seem consistent with existing WABC observations
(see Sect. ) and presumably reflect the strength of remote
forcing processes. In fall, poleward flow intensification has been related to
a major upwelling wind relaxation in the Gulf of Guinea. In contrast, Ps
flow intensification appears to be a more complex combination of local and
remote responses with time lags resulting in partial compensations.
WABC and Rossby wave dynamics
As presented above, the across-shore structure of the WABC and its seasonal
evolution are strongly suggestive of the important role played by westward
Rossby wave propagation. So is the seasonal evolution of the meridional flow
patterns that increasingly tilt away from the north–south axis in a clockwise
manner as they progress westward (Fig. ) owing to the rapid
change in Rossby wave phase speed in the tropics . More
precisely, Figs. , , and are
consistent with the generation of semiannual Rossby waves at the WA eastern
boundary via the scattering of coastal waves due to the meridional gradient
of the Coriolis parameter . Elements of linear
theory are recalled first as a starting point . Assuming no
meridional structure to the wave, i.e., the horizontal wavenumber is k=k,0, the Rossby wave dispersion relation is
ω=-βkk2+1/Rn2,
where Rn is the deformation radius for a given vertical mode n. Rn can
also be expressed as a function of the gravity wave speed cn for that
mode: Rn=cn/f. The (zonal) phase speed of a monochromatic wave of
frequency ω and zonal wavenumber k is thus
cϕ=ωk=-βk2+1/Rn2.
More complex perturbations propagate at the group velocity
cg=∂ω∂k=βk2-1/Rn2k2+1/Rn22.
Finally, for a given deformation radius and wave period there is a
critical latitude λn(ω) beyond which the free Rossby wave mode
is evanescent because the associated wavenumber solution to the quadratic
Eq. () has a nonzero imaginary part. At this critical
latitude f/β=cn/2ω so λn is defined by the fact
that tan(λn)=cn/(2ωRT) where RT is the Earth
radius.
In what follows baroclinic mode characteristics are determined based on the
TROP025 stratification computed on the model 75 grid levels. The calculation
is made with the dynmode program available at
https://woodshole.er.usgs.gov/operations/sea-mat/klinck-html/index.html
(last access: 18 August 2018). Using the gravity
phase speeds computed at 14∘ N and indicated in
Fig. , we find 22, 11, and 7∘ N for the critical
latitudes corresponding to the semiannual frequency and vertical modes 1, 2,
and 3, respectively. These estimates are in good agreement with those of
. The gravity phase speed of mode 3 is slightly higher
around 7∘ N than at 14∘ N
(0.85 m s-1 at 8∘ N vs. 0.70 at
14∘ N), so a more accurate estimation of λ3 is 9.5∘ N.
Parameters associated with each potential semiannual RW mode at 8
(baroclinic mode 1 to 3) and 14∘ N (baroclinic mode 1 and 2).
Gravity phase speed (m s-1), deformation radius (km), critical latitude (km),
theoretical wavelength (km), observed TROP025 wavelength (km), theoretical
phase speed (cm s-1), theoretical group velocity (cm s-1), and observed TROP025
propagation speed (cm s-1) are reported. No mode 1 RW is identified at
8∘ N. A mode 2 RW is also not identified at 14∘ N, in
agreement with linear theory predicting its evanescence. Note the important
discrepancies between theory and model RW mode 1 at 14∘ N.
cnRnθnλnλTROPcϕcgcTROP(m s-1)(km)∘ N(km)(km)(cm s-1)(cm s-1)(cm s-1)14∘ N, n=12.1592211206507.05.63.514∘ N, n=21.13212–––––8∘ N, n=12.2112224080–2624–8∘ N, n=21.36412138013108.57.17.48∘ N, n=30.85429.54835503.01.63.0
Wavenumber–frequency dispersion diagram for baroclinic Rossby waves
corresponding to values of the deformation radius equal to 112, 64, and 42 km
(solid black, baroclinic modes 1–3 at 8∘ N) and 59 and 31 km
(solid gray, baroclinic modes 1 and 2 at 14∘ N). The black (gray) filled dots and
dashed lines represent the Rossby wave characteristics
identified in Fig. for TROP025: ω=2π/(6 months);
kTROP=2π/λTROP and cTROP as listed in
Table . The gray dash-dotted line is the dispersion curve
that passes through the Rossby wave mode identified at 14∘ N in
TROP025. Its associated deformation radius is 48 km, i.e., significantly
less than the actual baroclinic mode 1 deformation radius at that latitude
(59 km). The horizontal dotted line corresponds to the semiannual
frequency.
Our aim is now to describe the semiannual Rossby wave dynamics present in
TROP025, assess the degree to which the simplest linear theory captures this
dynamics, and gain insight into the different baroclinic mode contributions.
Dispersion diagrams are shown in Fig. for the Rossby wave
modes permitted at 8 and 14∘ N, modes 1 to 3 for the former and mode
1 for the latter. The dispersion curve for mode 2 at 14∘ N is also
shown. The fact that it does not intersect the horizontal line corresponding
to semiannual frequencies underscores the evanescent nature of mode 2 RWs at
this latitude. Model wavenumbers and propagation speeds deduced from careful
examination of Fig. are reported in Fig. . At
8∘ N a slow and a fast wave coexist. They fall quite accurately on
mode 2 and 3 dispersion curves, respectively. On the other hand, the dominant
wave identified at 14∘ N is distinct from a linear mode 1 RW (see
also Table ).
To elaborate on the significance and relative importance of the different RW
modes, a harmonic analysis was performed at each grid cell (xi,yj,zk)
over the period 1982–2012 to extract the semiannual variability of the
meridional velocities. The resulting semiannual harmonics (6-monthly values)
were subsequently decomposed onto the baroclinic modes computed for each
location (xi,yj) based on a local annual mean profile of Brunt–Vaisala
frequency. Horizontal maps of the depth-integrated kinetic energy associated
with modes 1–4 are shown in Fig. , after time averaging
over the semiannual cycle. Restricting the final step of vertical
integration to the layer in which the poleward flow is concentrated, e.g., the upper 200 m,
leads to similar results and conclusions. The dominance
of mode 2 is a well-known attribute of many equatorial and tropical regions
. Mode 2 dominates over most of the ETNA except for a few
isolated offshore grid cells in the latitude range 10–15∘ N where
mode 1 is of comparable magnitude. The shape of the region having finite
values of mode 2 kinetic energy and the general offshore decay seen in
Fig. b are consistent with the following: energy being mainly radiated
from the coastal waveguide where the largest energy values are found; and westward
energy propagation being more effective at lower latitudes and ineffective
north of 12–15∘ N, with a noticeable change in the cross-shore size of
the region with finite energy values around 10∘ N. A similar impression can be
drawn for mode 3 and 4 except that westward propagation seems both more
strongly damped and more confined meridionally in a low-latitude band. This
is qualitatively consistent with the λn values decreasing with mode
order (Table ). The kinetic energy distribution for mode 1
is peculiar and does not exhibit large values in the coastal waveguide
(Fig. a). Why so little of the semiannual CTW activity
projects onto baroclinic mode 1 is a pending question left for a future
investigation.
Vertically integrated kinetic energy (m3 s-2) for the
semiannual meridional velocity decomposed onto baroclinic modes. Only the
first four modes are shown. Time averaging is performed over the semiannual
cycle. The two white dashed lines represent the location where zero PV
gradients are found in Fig. .
Vertical structure of the ETNA baroclinic modes 1 (blue), 2
(green), and 3 (red) for pressure and horizontal velocities (upper 2500 m
only). Calculation is made using TROP025 stratification at 14∘ N,
18∘30′ W. The associated reduced gravity phase speed ce
( s-1)
for each mode is also indicated.
Westward propagation of energy away from the coastal guide is an important
process contributing to the poleward attenuation of the WABC. Over the
continental slope, the progressive deepening of the WABC with increasing
latitude (Fig. ) corresponds to a reduction (increase) in the
relative contribution of high-order (low) modes, e.g., less weight on mode 3 whose upper zero crossing is at 75 m
(Fig. ). Although this is consistent with idealized
simulations and theoretical arguments , we
cannot be certain that TROP025 motions associated with high-order modes along
WA are more efficiently damped for physical reasons (such as RW generation
and frictional processes) as opposed to being dissipated by excessive
numerical viscosity and/or diffusion. For a given mode n dissipation of numerical
origin should increase as latitude increases and the corresponding typical
horizontal scale associated with that mode Rn decreases. The realism of the
model WABC may thus deteriorate with increasing latitude.
In this context, the model behavior in the latitude range ∼14–22∘ N requires further clarification. The northern end of this
sector coincides with the critical latitude for baroclinic mode 1 semiannual
RWs; hence the possibility of resonant excitation of these waves because
their group velocity vanishes . On the other hand,
Hovmüller diagrams similar to Fig. for latitudes between 15
and 22∘ N reveal a dramatic reduction of the semiannual RW signal
toward the north (not shown but see Figs. and
for indirect evidence). This latitude range corresponds to
a major transition in the Canary Current system with distinct WSC forcing
conditions and dynamical regime on either side (offshore conditions
associated with negative wind stress curl and equatorward flow prevail north
of 20∘ N), two abrupt geomorphological near discontinuities (at Cape
Verde and Cap Blanc), and the permanent Cape Verde thermohaline frontal
zone. All these sources of nonlinearities can contribute to the northward
weakening of the semiannual CTW signal and thus prevent the generation of
semiannual RW activity beyond 18–20∘. They can also explain the
discrepancies found at 14∘ N between RW characteristics expected
from linear theory and those identified in the model.
However, the upper ocean potential vorticity (PV; see Sect. )
field offers another compelling explanation for the meridional structure of
the RW field found in TROP025. Equation () is strictly valid in a
large-scale ocean at rest in which the only source of background PV gradient
is the Coriolis parameter gradient df/dy=β. In realistic conditions,
the large-scale PV field implicated in the propagation of baroclinic Rossby
waves must account for stretching effects associated with background shear
flows if any . More appropriate quantities to
investigate upper ocean RW dynamics are total PV gradients in three density
layers (see definition in Sect. ): 25.2–26.3 (layer 1), 26.3–26.7
(layer 2), and 26.7–26.9 (layer 3). Layers 1 and 2 are layers in which a large
fraction of the WABC transport is concentrated (Fig. ).
They are of comparable thickness and typically occupy the upper 200–250 m.
Layer 3 is also of comparable thickness but it is associated with a modest
fraction of the poleward transport, both nearshore (Fig. )
and offshore (Fig. ). PV fields calculated following
Eq. () are shown in Fig. as is their
gradient vector field. Layers 1 and 2 exhibit relatively similar patterns. PV
gradients in these layers strongly depart from those resulting from
variations in the Coriolis parameter alone. In particular, a reversal of the
gradient is found along oblique lines that run
northeast to southwest between Cap Blanc and the Cape Verde islands. RWs approaching these lines
must be subjected to intense dispersive, refractive, and/or dissipative effects. Layer
3 is the deepest layer where PV gradients are not uniformly oriented toward
the north (the gradients vanish in a broad northern sector where
stratification and Coriolis parameter contributions nearly cancel). Below
layer 3 PV gradients are dominated by β and relatively uniform from the
coast to thousands of kilometers offshore (not shown).
Shallow-water potential vorticity field in the ETNA computed from
TROP025 (color; 10-8 m-1 s-1) using Eq. () for the
density classes 25.2–26.3 (a), 26.3–26.7 (b), and 26.7–26.9 (c). These values are
such that the three layers are of comparable thickness and the same
color scale can be used. Potential vorticity gradients are also shown in
vectors. Note the vanishing gradients found along the white dashed lines in
panels (a) and (b) (the position of the two lines differ slightly, e.g.,
see their position with respect to the Cape Verde islands).
At 14∘ N, the zero PV gradient line in layer 1 (layer 2) is
located at ∼24∘ W (28∘ W) so the distance to the WA shelf
break is not enough (marginally enough) to fit a linear RW wavelength
(λ1=1120 km; see Table ) in the sector where PV
gradients are mainly directed from south to north and relatively uniform
spatially. We suspect that the differences in PV gradients between layers 1–2
and the deeper layers are implicated in the deviations from standard linear
theory (wavelength and propagation speed) and also in the rapid RW signal
attenuation observed within 1000 km from shore (Figs. and
) despite the critical latitude being located far to the
north. Importantly, the width of the sector situated east of the zero PV
gradient decreases rapidly with latitude between 14 and 20∘ N and is
only ∼300 km wide at 18∘ N. This makes the existence of long
weakly dispersive semiannual RWs increasingly implausible when approaching
Cap Blanc.
Overall, the modifications of the ETNA PV field by vortex stretching effects
in the density range 25.2–26.7, at which most of the meridional flow and Rossby
wave energy are concentrated, appears as a good candidate to explain the
(meridionally variable) cross-shore damping scale for RWs and the progressive
reduction of RW amplitude north of 15∘ N. At deeper depths, the
impact of the PV gradients in the upper ETNA may be more limited.
Across-shore sections reveal alternating bands of poleward and meridional
flow below 300–400 m that migrate westward past the zero PV gradient lines identified above (not shown), suggesting the presence of Rossby
wave activity there in agreement with the observational findings of
. The physical processes responsible for the particular PV
structure present in the upper ETNA will be discussed in Part 2.
Excitation of free Rossby waves by wind stress curl forcing along WA can also
contribute to RW activity and in turn impact the WABC seasonal cycle. This
is expected to arise if the WSC spatiotemporal patterns of variability in
the immediate vicinity of WA involve particular wavenumber–frequency pairs
consistent with the dispersion relation of free RWs . South
of 20∘ N, the wind stress curl and Sverdrup transport fields shown in
Fig. exhibit across-shore spatial variations with a
contribution of wavelengths 500–1500 km. This is in the appropriate range
to excite baroclinic mode 1 and 2 RWs south of their critical latitude. But
semiannual variability of the wind curl signal is particularly weak in the
WA sector (see Fig. ) so its contribution to the WABC
semiannual cycle must be limited. Based on frequency considerations the
generation of an annual RW signal is more plausible. However, a harmonic
analysis similar to the one described above reveals small amounts of energy
associated with the annual cycle of the upper ocean meridional flow (not
shown). We relate this to a combination of inhibiting factors. South of
15∘ N mode 1 RWs have wavelengths ∼ 2500 km or more, i.e., too large to be compatible with the WSC typical spatial scales of
variability. North of 10–12∘ N such RWs are also too large to fit in
the sector situated east of the line where PV gradients vanish.
Discussion and conclusions
An eddy-permitting numerical simulation with realistic forcings has been
analyzed to investigate the dynamics of the boundary current along the West
African seaboard. The depth range of interest was chosen to be above the
σt=26.7 isopycnal. This broadly coincides with the upper 250 m of
the water column and places the focus on the layer of fluid in which the
wind-driven circulation is overwhelmingly concentrated. The geographical
focus is roughly on the southern sector of the Canary Current system between
∼8 and 20∘ N. In this area wind stress curl (both nearshore
and offshore) is robustly positive, i.e., conducive to poleward
flow
This is in contrast to most other eastern boundary upwelling
systems in which offshore and nearshore wind stress curl tend to be of opposite
sign .
. In fact, upper ocean equatorward currents are rarely
found over the WA continental slope, except when approaching 8–10∘ N
where WSC is much reduced. The model poleward flow is characterized by two
main intensification periods in spring and fall. We interpret this
characteristic as a consequence of the low-frequency coastal trapped wave
activity generated locally and remotely by seasonal wind fluctuations along
the African shores
In contradiction to the assertions made in
various places including we were unable to establish
a connection in the model between the summertime pulse of the NECC and the
fall near-surface intensification of the WABC. Our unsuccessful attempts
included diagnostics aimed at tracking the propagation and advection of patterns
of elevated surface pressure signals from the region 23∘ W, 8–10∘ N
(where the northern NECC summer pulse is strongly felt; Fig. )
toward the east–northeast where they could contribute to enhancing alongshore
pressure gradients at the WA continental slope. In Part 2 we will show that
timescales associated with advection are too long for this to happen (Rossby
waves propagate pressure signals toward the west and are therefore not
candidates).
. This is an important difference from the annual cycle of the
boundary slope current discussed in several past studies including
and . Note, though, that important
signs of a semiannual cycle can be seen in in which two
along-slope transport maxima are found across their so-called section B.
Despite some differences in their forcing regions and precise depth–latitude
range of influence, the spring and fall model WABC intensifications bear
important similarities. They are both part of a semiannual cycle whose
forcing processes were carefully investigated. We found no clear signs that
along-slope motions in the form of free Rossby wave modes could be excited by
nearshore WSC, perhaps because the WSC temporal variability at semiannual
frequency is particularly small. Conversely, along-slope flow generation and
propagation in the coastal waveguide is prominent as we found throughout the
study. Considered in isolation, the WABC spring intensification accompanies
the relaxation of wintertime coastal upwelling winds in the latitude range
7–20∘ N as the ITCZ shifts northward toward that area. Flow
intensification is found in the subsurface and is broadly consistent with the
theory on undercurrents in upwelling systems. Time lags between the
contributions of the wind for different sectors along the African coast are
such that compensating effects occur and the WABC forcing is suboptimal. The
fall intensification is stronger. It has a more remote and focused origin
that we are able to locate in the Gulf of Guinea through spatiotemporal
analyses of both wind forcing and coastal ocean response. Owing to this
remote forcing, the largest WABC transports occur while WSC is relatively weak
and upwelling winds intensify, i.e., when local forcing is least
conducive to poleward currents. In that respect, our work tends to
substantiate old assertions about the connection between the boundary current
flowing offshore of Senegal and Mauritania as well as poleward flow in the Gulf of Guinea,
albeit only during part of the annual cycle. Conversely and in contrast to
what has been hypothesized for the southeast tropical Atlantic
, wind variability and Kelvin wave activity in
the equatorial Atlantic are not found to be implicated in the forcing of the
WABC semiannual cycle . Our results also differ from
those for the eastern South Atlantic in that the seasonal
cycle of the WABC is not directly linked with the local wind stress curl,
which as mentioned above has very little semiannual modulation.
More quantitatively, the model provides estimates for poleward transport
over the WA continental slope. They depend somewhat on the precise choices
made for the control surface (depth, across-shore integration bounds, and
position in latitude). At the latitude of Senegal (Mauritania),
geostrophic transport above σt=26.7 is of the order of 1 Sv (0.4 Sv), i.e., a large (moderate) fraction of the theoretical
barotropic Sverdrup transport. We relate this to the meridional changes in
WABC dispersion through Rossby wave generation. Indeed, dispersion is most
pronounced at low latitudes where Rossby waves travel faster and higher
baroclinic modes can be impacted. The vertical structure of the boundary
currents reflects these differences. Upper ocean confinement of the Sverdrup
flow by Rossby waves is systematically most
pronounced at lower latitudes; hence the north–south differences in depth
range of the WA boundary current. Transport distribution at 8∘ N is
also consistent with these assertions but WSC diminishes greatly when
approaching such low latitudes where our analyses may be more difficult to
interpret.
The model analysis of Rossby wave activity reveals important differences from
previous descriptions of these waves at larger scale in the
eastern to central North Atlantic . Most notably, we find that
upper ocean Rossby wave activity generated at the eastern boundary remains
confined to a well-defined ocean sector delimited by the WA seaboard and an
oblique line running northeast–southwest where background potential
vorticity gradients vanish in the upper ocean. Such modification of the
potential vorticity field is due to stratification effects having a magnitude
comparable to planetary vorticity effects (β). Past studies of Rossby
wave activity in the North Atlantic have classically been made using linear 1.5 layer reduced gravity models
with no background flow that RWs can interact with. It seems to be an
important limitation for the ETNA sector situated north of about
12∘ N.
Our main qualitative conclusions on the processes responsible for the WABC
semiannual cycle and its associated RW dynamics in our model are
schematically summarized in Fig. . The relevance of this
numerical investigation to the real WA ocean is an obvious concern. Although
TROP025 skills at regional and basin scale have been demonstrated
(Sect. ) model biases cannot be excluded at the scale of the WABC.
Because there have been relatively few observational programs in this part of
the world ocean we can only offer limited and qualitative insight into model
realism, for example on the reality of the two WABC intensification phases
and the associated flow characteristics. Although the existence of two
poleward intensification phases is not systematically recognized in previous
studies, published observations are not inconsistent with the model behavior,
including on the timing of these two phases.
Schematic representation of the main processes relevant for the
WABC seasonal cycle as identified in this study. Two WABC intensification
phases occur in spring (S) and fall (F) in relation to the relaxation of
cross-shore Ekman transport along different sectors of the African coast. The
relative importance of each sector varies depending on the season. The
font size of the S and F letters next to the WABC arrows is indicative of this
importance. Three sectors are distinguished: the Gulf of Guinea (GG), the
West African corner (WAC), and West Africa (WA). The fall intensification is
mainly due to wind changes in the GG and a modest contribution from the WAC.
The spring intensification is the outcome of a subtle combination of forcings
in all three sectors with a prominent role of local WA forcings. The
spatiotemporal variability of the wind stress curl forcing does not have an
appreciable effect on the WABC seasonal cycle, although it is a key local
determinant of the climatological mean meridional transport along WA (weak
near 8–10∘ N and particularly strong near 20∘ N). The WABC
is subjected to scattering and gives rise to Rossby wave (RW) activity
hundreds to thousands of kilometers offshore. The characteristics of these
RWs exhibit rapid meridional changes over a few degrees of latitude. At low
latitude (∼8) where β is largest RWs associated with baroclinic
modes 2 and 3 dominate. At 14∘ N where only mode 1 RWs are permitted
the model RWs significantly differ from those predicted by standard linear
theory. This discrepancy and the progressive disappearance of RW activity
north of 15∘ N may be related to the peculiar structure of the upper
ocean ETNA PV field: above approx. 250 m of depth the gradient of the
background PV field (β+ stratification effects) vanishes along a
line running northeast to southwest.
The field experiment CANOA08 took place in November 2008 at a time of year
when the poleward flow should be most intense including near the surface.
Above σt=26.85 the flow over the continental slope carried SACW up
to 25∘ N where vanishing meridional transports were
found, i.e., exactly the November climatological limit determined for
the same density class in the model (not shown). In addition, observed
transports are broadly consistent with those found in TROP025, particularly
for the uppermost stratum examined by (see model–data
comparison in Table ). In CANOA08 the intense poleward
surface and subsurface currents found in the vicinity of Cap Blanc at
surface and subsurface levels are interpreted by the authors as,
respectively, “a late expression of the summer Mauritania current” and a
local response to strong upwelling winds . Our model results
cast doubt on these interpretations and suggest instead that CANOA08 may
have sampled the ETNA at the (normal) time when the fall intensification of
the surface and subsurface WABC is remotely forced.
Meridional transports observed in November 2008 reported in (left) and their TROP025
climatological equivalent for the month of November (right) separated by the / symbol. Transport values (in Sv) are
provided for two different density layers (SW for surface waters, σt<26.46; uCW for upper
central waters, 26.46<σt<26.85) and four latitudes (16.25, 17.5, 20, and 24∘ W). In situ
values are estimated visually from Fig. 7a–c in for the cross-sectional area between the
shelf break to 60 km offshore. For the model, transport is computed over four grid cells (100 km) situated
offshore of the shelf break. Using this wider across-shore section is meant to account for our limited
horizontal resolution, but transports estimated over only three grid cells only differ by 15 %–20 %.
On the other hand, comparison with CANOA08 observations raises some concerns
about the along-slope continuity or coherency of water mass transport in
TROP025. In the model, the low-salinity signal characteristic of SACW does
not penetrate north of Cap Blanc irrespective of the season
(Fig. ). This is inconsistent with observations reported in
(as well as older ones) showing large amounts of low-salinity
SACW up to 24∘ N. The across-shore exchanges of water between the
WABC core and the open ocean may thus be overestimated in the model,
plausibly as a consequence of insufficient horizontal resolution. Given the
difficulty of maintaining moorings in this area multiyear repeats of the CANOA
array at different seasons would provide useful information on the temporal
variability of the WABC.
To our knowledge, the only data that are available to assess model realism
during the spring WABC intensification are those obtained at
21∘40′ N as part of the JOINT-1 experiment to investigate shelf
and slope currents in the vicinity of Cap Blanc. Current meter data
their Figs. 4 and 6 underscore the importance of
synoptic variability with dramatic fluctuations of the poleward flow on timescales of days to weeks. Consequently, JOINT-1 efforts remain inconclusive
with respect to the mean structure of the boundary currents at that time of
year (including on the possible existence of a near-surface poleward
countercurrent over the continental slope). A similar issue plagues the study
of the drift of a parachute drogue released off Cap Blanc at 50 m of depth
. The trajectory reveals intense northward flow
(24 cm s-1 over a 12 h period) that cannot be considered
representative of average conditions. However, combined with the CINECA
February–March 1972 velocity observations at 19–20∘ N his
Fig. 9 showing a core of poleward velocities in excess of
8–10 cm s-1 in the depth range 100–200 m, the general impression is
that TROP025 underestimates the intensity of the flow at these latitudes
(based on Fig. and also on the examination of a figure
similar to Fig. for the latitude band 19–21∘ N; not
shown). This would be another possible reason why the low-salinity signal
associated with the presence of SACW north of Cap Blanc is not reproduced
in TROP025.
Overall, the realism of the model boundary circulation is uncertain given the
scarcity of available observations. In addition, our dynamical
interpretations frequently invoke baroclinic mode decomposition, which is not
strictly valid in the horizontally heterogeneous conditions under which we use it.
More elaborate approaches such as WKB ray tracing may prove useful in this
regard, e.g., to clarify the reasons why sea level and upper ocean
flow signals propagate offshore at different speeds. In this context, the
present study should be seen as a way to stimulate and guide future work in
this highly undersampled part of the world ocean. Part 2, which aims to
connect the WABC to the regional circulation context of the subtropical North
Atlantic shadow zone, shares the same general objective.
TROP025 model simulations are stored at the CEA TGCC
supercomputing center and will be made available upon request to the authors.
LK and XC designed the study. JJ performed the numerical simulation.
LK carried out most of the analyses, with contributions from JJ, XC, and NK; XC and LK wrote the manuscript with input from all authors.
The authors declare that they have no conflict of
interest.
Acknowledgements
The ART PhD program of the Institut de Recherche pour le Développement and
Lala Kounta was funded through the ART PhD program of the Institut de Recherche
pour le Développement while conducting most of this research. We also
acknowledge support from the AMMA2050 project (funded under the Future
Climate for Africa program by the NERC and DFID). Numerical simulations for
this work were performed through computing allocations GENCI GEN1140 on
Curie. We thank the editor and two anonymous reviewers for their careful
reading of and comments on the manuscript. We also thank Alain Colin de Verdière, Julie
Deshayes, Bernard LeCann, Richard Schopp, Jérôme Sirven, and Jérôme Gourrion for
useful suggestions and comments.
Edited by: John M.
Huthnance Reviewed by: two anonymous referees
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