Coastal upwelling dynamics

Ekman transport and pumping

According to Ekman’s theory (Ekman, 1905), low-level winds blowing above the ocean surface force a circulation below the surface, which is deviated to the right in the northern hemisphere and to the left in the southern hemisphere due to the Coriolis force. At the surface, the resultant water motion is at 45° from wind direction, clockwise (resp. counter- clockwise) in the northern (resp. southern) hemisphere. The turbulent tension associated to such motion forces a circulation in the water layer just below the surface layer, with an angle that depends on the coefficient of vertical eddy viscosity. The same process is repeated for each layer of the water column, with an exponential decay of the associated currents.

Horizontal motion associated to surface wind forcing can be described over the water column from the surface to the bottom by the Ekman spiral (fig. 1.1). By integrating the associated equations over the vertical, one obtains the net zonal and meridional transports Tx and Ty:

T_x f^y ρτ

= T_y f^x

ρτ

−

=

where τx and τy are for the zonal and meridional components of the surface wind stress, ρ is for seawater density and f is for the Coriolis parameter:

ϕ sin 2Ω

= f

where Ω is for the earth’s rotation rate and φ is for latitude.

Fig. 1.1: Schematic of the Ekman spiral and transport in the northern hemisphere.

Hence the net transport of water called the Ekman transport is at right angles with the wind direction, and is directed to the right (resp. to the left) of wind direction in the northern (resp.

southern) hemisphere, as indicated on figure 1.1. Because of the exponential decay of water motion, the Ekman transport is mainly confined in a top layer called the Ekman layer. Below the Ekman layer, the effect of the wind forcing on the ocean currents is weak and can be easily neglected.

Fig. 1.2: Schematic of coastal upwelling in the HCS. The pink arrow is for the alongshore winds, and the blue arrow is for the offshore Ekman transport and the associated upwelling.

In the case of EBUS’es, the Ekman transport is of fundamental importance for the understanding of ocean dynamics. Consider the case of a quasi-constant wind blowing parallel to the coast with an equatorward direction, as it is generally the case for EBUS’es. The resulting Ekman transport will thus have a westward direction and warm waters of the Ekman layer move away from the boundary in the offshore direction (fig. 1.2). To compensate this

loss of volume, cooler waters from deeper layers are advected shoreward and upward when they reach the boundary: this phenomenon is called coastal upwelling (fig. 1.4).

There is a secondary mechanism that also contributes to coastal upwelling: Ekman pumping. Ekman pumping is due to spatial variations of the surface winds and of the associated Ekman transports, that induce areas of convergence and areas of divergence of the waters in the Ekman layer. In areas of divergence (resp. convergence), the conservation of water volume drives upward (resp. downward) vertical velocities at the base of the Ekman layer, ie upwelling (resp. downwelling). These velocities are called the Ekman pumping (WEk), which can be written:

⎟⎟⎠⎞

⎜⎜⎝⎛

∂∂

∂ −

= ∂

y x W_Ek f τ^y τ^x

ρ 1

(a) (b)

Fig. 1.3: Schematic of the effect of Ekman Pumping on a southern hemisphere EBUS. The pink arrows are for the alongshore winds, the dark blue arrows are for the offshore Ekman transports and the light blue arrows are for the associated upwelling (a) or downwelling (b).

In the presence of a coastline, which is the case for EBUS’es, the alongshore winds are stronger over the ocean than over the land because the friction at the surface is increased over the continent due to the presence of orography and vegetation which tend to reduce the wind intensity. This creates a cross-shore gradient of alongshore wind intensity, which is reduced near the coast in the so-called wind drop-off zone. As explained by Bakun and Nelson (1991), the resultant wind stress curl drives upward Ekman pumping which contributes to the coastal upwelling, together with Ekman transport. This mechanism is schematized on figure 1.3 (a).

Likewise, if the wind were stronger near the coast, it would induce stronger Ekman transport near the coast than farther offshore: an area of convergence would appear off the coast and the

Ekman pumping would be negative and have a downwelling effect that counteracts the upwelling effect of Ekman transport (fig. 1.3 b).

Coastal Currents

Fig. 1.4: Vertical section in the cross-shore direction of the mean alongshore currents (m/s) and the mean temperature (°C) simulated by the ROMS model (Shchepetkin and McWilliams, 2005) at 10°S near the central Peru coast. Colours and black contours are for currents. Full (resp. dashed) lines are for positive (resp. negative) values, ie for equatorward (resp. poleward) flow. White contours are for temperature (interval is 1°C). The x-axis (resp. y-axis) is for cross-shore distance (resp. depth) and the corresponding units are km (resp. metres).

A similar structure of coastal currents constituted by an equatorward surface current and a poleward subsurface current (cf. figure 1.4 for the HCS) is observed for all EBUS’es.

The mechanism controlling the equatorward surface current is well-known: coastal upwelling causes a rise of the isotherms and isopycns near the coast due to the cooler and denser waters upwelled at the surface (fig. 1.4). Hence, a cross-shore density gradient is present in the coastal zone, with denser waters near the coast and less dense waters offshore: this induces an equatorward flow close to the coast due to geostrophy. In addition, direct wind forcing in the Ekman layer also drives an equatorward flow (fig. 1.1), which combines with the geostrophic flow to form the surface current. On the other hand, so far there is no consensus on the origin of the poleward undercurrent observed along EBUS’es. One theory relates it to the alongshore meridional pressure gradient: waters are increasingly warmer towards the equator, causing the sea level to be higher which creates a poleward flow (Neshyba et al., 1989). Another theory based on the results of a shallow-water baroclinic ocean model in the idealized case of a

meridional eastern boundary and spatially-uniform alongshore wind forcing relates both the equatorward surface current and the poleward undercurrent to the different baroclinic modes and succeeds in reproducing the depth and magnitude of the observed currents (McCreary, 1981). The poleward undercurrent in particular is related to the effect of bottom friction over the continental slope, which damps the higher order modes and thereby enhances the control of the undercurrent by the intermediate order modes (McCreary and Chao, 1985).

Mesoscale and sub-mesoscale processes

Fig. 1.5: Meanders of the upwelling front indicated on a map of SLA in the southern HCS simulated by the ROMS model (1/6° resolution, open boundary conditions from a global ocean model, atmospheric forcing from observed satellite winds and reanalyzed air/sea fluxes) for 19^th August 1995. Unit is meter. Contour interval is 0.02 m. The black, red, green and blue dashed lines indicate the position of the (-6 cm)-isoline on 19^th August, 24^th August, 29^th August, and 3^rd September 1995, respectively. White ellipses indicate the position of a few meanders (V. Echevin, personal communication).

Coastal upwelling induces a strong cross-shore thermal gradient in the nearshore area, with cooler waters upwelled near the coast and warmer surface waters pushed away from the coast by the combined effects of Ekman transport and Ekman pumping. As a consequence, a sharp temperature front called the upwelling front is present a few dozens of kilometres offshore. The position and intensity of this front is subject to wind variability, which can be rather strong on intraseasonal time scales (e.g. Garreaud and Muñoz, 2005; Renault et al., 2009). In addition, upwelling-favorable alongshore winds are also subject to significant

alongshore variability due to the varying continental orography and the shape of the coastline:

for instance, winds are generally locally stronger around capes because the land-sea surface occupation ratio is smaller, and weaker around bay areas. The combined spatial and temporal wind variabilities create small scale perturbations along the upwelling front, which can take various forms, namely meanders, filaments (also called plumes) and eddies. Meanders are mesoscale and sub-mesoscale (typical length is a few dozen kilometres) large-curvature portions of either the upwelling front or the coastal currents (fig. 1.5). Filaments are sub- mesoscale intrusions of the upwelling front into the offshore region (fig. 1.6), and are characterized by high primary production and biological activity.

Fig. 1.6: Sub-mesoscale filaments indicated on a map of SST in the eastern South Pacific from AVHRR Pathfinder satellite data (Vazquez et al., 1995) for 1^st January 1992. Unit is °C. Contour interval is 1°C.

Eddies, which can be cyclonic – counter-clockwise (resp. clockwise) geostrophic currents in the northern (resp. southern) hemisphere – or anticyclonic – clockwise (resp.

counter-clockwise) geostrophic currents in the northern (resp. southern) hemisphere – are mesoscale features with a radius of the order of 100km (fig. 1.7) and are common in EBUS’es (Chaigneau et al., 2009). Cyclonic (resp. anticyclonic) eddies are also called warm-core (resp.

cold-core) eddies because they are associated to a positive (resp. negative) sea level anomaly (SLA) and a warm (resp. cold) sea surface temperature (SST) anomaly, as illustrated on figure

Filaments

1.7. Like the smaller scale meanders and filaments, they can be generated by spatio-temporal variations of the upwelling front, which are the result of baroclinic instabilities generated by vertical shear between the equatorward surface current and the poleward undercurrent (e.g.

Leth and Shaffer, 2001; Marchesiello et al., 2003). Mesoscale eddies are generated near the coast and propagate westwards: for this reason, they are important contributors for the transfer of physical properties from the coastal region to the open ocean. Their nearshore spatial variability can also determine the location of spawning areas for small pelagic fish species (Logerwell et al., 2001).

Fig. 1.7: Cyclonic and anticyclonic mesoscale eddies indicated on a map of SLA in the eastern South Pacific simulated by the ROMS model for 1^st January 1992. Unit is meter.

Coastal-trapped waves

The presence of poleward-propagating coastal-trapped waves (CTW) has been reported along the west coast of North and South America (e.g. Brink, 1982; Chapman, 1987) and along the west coast of Africa (Polo et al., 2008). Such baroclinic waves are forced by equatorial Kelvin waves (EKW) travelling eastward across the equatorial Pacific and Atlantic as they impinge on the eastern boundary (Clarke, 1983; Enfield, 1987): according to theory

Anticyclonic eddies Cyclonic eddies

(Clarke, 1983; Grimshaw and Allen, 1988; Clarke and Shi, 1991), for each baroclinic mode associated to an EKW of any given frequency, a critical latitude can be found that separates the zone of CTW propagation along the ocean boundary (poleward of that latitude) from the zone of reflection of the EKW as westward-propagating equatorial Rossby waves (ERW) and extra-tropical Rossby waves (ETRW) (equatorward of that latitude). CTW are associated to anomalies of sea level and velocity, which can destabilize the thermocline and the coastal currents (e.g. Shaffer et al., 1997; Pizarro et al., 2002). Such perturbations are thought to have a direct impact on the behaviour of the marine ecosystem in EBUS’es, as proposed by Bertrand et al. (2008) for the HCS.

The four main EBUS’es share common characteristics that were briefly described above: upwelling-favorable alongshore winds that drive intense primary production and biological activity through physical processes of Ekman transport and pumping; alongshore coastal currents with equatorward flow at sea surface and poleward flow below the surface;

intense mesoscale and sub-mesoscale activity; poleward-propagating coastal-trapped waves.

However, each EBUS also has its own specificities, related to wind regime, coastline orientation, bottom topography, continental orography, latitude, among other factors. What are the peculiarities of the HCS compared to other EBUS’es ? What makes it potentially more sensitive to climate change than the other systems? This is the object of the next section.

1.1.2. The Humboldt Current System: a highly productive ecosystem under