Coastal currents in low latitudes (with application to the Somali and El Niño currents)

Deep-SeaResearch,Vol.30, No.8A,pp. 887to 902,1983. PrintedinGreatBritain.

0198-0149/83$03.00+ 0.00 PergamonPressLtd.

Coastal currents in low latitudes (with application to the Somali and El Nifio currents) S. G. H. PHILANDER* and P. DELECLUSE*t

(Received25 May 1982; accepted27 September 1982) Abstract--Near the equator, the directly wind-drivencoastal jet along the western boundary of an ocean basin differs from its counterpart along the eastern boundary in several respects. (1) The western boundary current is more intense by an order of magnitude, be~anse it flows in the direction of the pressure force associated with density gradients whereas the eastern boundary current is opposed by this pressure force. (2) In the west the flow at the depth of the thermocline is still in the direction of the wind, but in the east the pressure force drives an undercurrent that is practically as strong as the surface current. (3) Variability on time scales from a week to months is much higher in the east than the west. (4) On time scales longer than a month long Rossby waves disperse the eastern boundary current westward, but the western current remains a coastal jet. (5) A relaxation of the wind causes a prompt reversal in the direction of the flow in the east but in the western region, where the current flows 'downhill', there is only a gradual deceleration. The relevance of these results to the Somali Current in the western Indian Ocean and El Nifio Current in the eastern equatorial Pacific Ocean is discussed.

INTRODUCTION WINDS parallel to coasts drive coastal currents and cause upwelling at all latitudes, but near the equator the coastal response is distinctive. In the western Indian Ocean for example, the Somali Current, a coastal jet driven by the southwest m o n s o o n s parallel to the east African coast, is discontinuous and has two branches. The southern branch veers offshore in the neighborhood of 4 ° N ; the northern branch is part of a clockwise gyre between 4 and 1 0 ° N (Fig. 1). There is nothing special about the structure of the winds near 4 ° N , but upwelling is most intense in this separation region, which is marked by a tongue of cold surface waters (DOING, MOLINARI and SWALLOW, 1980; BROWN, BRUCE and EVANS, 1980). Along the western coasts o f equatorial Africa and South America the winds are also parallel to the coast. Here we expect narrow coastal currents on short time scales only, because on long time scales Rossby waves disperse the currents westward so that they become broad and slow (PHILANDER and PACANOWSKI, 198 l). This is presumably the reason why the narrow, warm, southward-flowing El Nifio:~ current along the coasts of Ecuador and northern Peru persists for at most a month or two, F e b r u a r y and March (SVERDRUP, JOHNSON and FLEMING, 1942). There is a similar current along the western coast of equatorial Africa

* Geophysical Fluid Dynamics Laboratory/NOAA, Princeton University, Princeton, NJ 08540, U.S.A. t Present address: Laboratoire d'Oceanographie Physique, Museum National d'Histoire Naturelle, Paris, France. :~The term El Nifio is currently also used for interannual phenomena characterized by warm sea-surface temperatures over much of the tropical Pacific Ocean. 887

888

S.G. tt. PHILANDERand P, DELECLUSE

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Fig. 1. The distribution o f surface currents between 6 June and 30 July 1979 o f f the eastern coast of Africa. Current arrows are c e n t e r ~ on the point o f observation. After DOING, MOLINARIand SWALLOW(1980).

personal communication). It is curious that the low latitude coastal currents along the eastern boundaries of the ocean baskss appear when the winds arc very weak and then flow in a direction opposite to that of the wind. Why are the eastern boundary currents so different from the Somali Current and why are the coastal currents in low latitudes so different from those in higher latitudes? Far from the equator, coastal currents evolve in two stascs. Immediately after the sudden onset of spatially uniform winds parallel to the coast, alongshore variations arc ncgl/sihlc. During this phase the wind accclczatcs the coastal jet, whose width is the radius of deformation, and offshore Ekman drift induces coastal upwell/n$ (CaAR~Y, 1955). Coastal Kelvin waves, excited at the boundary of the forced region, establish alongshore variations within a radius of deformation of the coast and change the flow dramatically by introducing an alongshore pressure gradient that balances the wind stress. This causes the jet to stop accelerating. Horizontal divergence of the jet now maintains the offshore Ekman drift so that the upwelling is sharply reduced after the passage of the Kelvin wave (ALLEN, 1976). The neighborhood of the equator differs from higher latitudes because there, meridional gradients are large before the arrival of any Kelvin waves. This is so because a meridional

(P. H1SARD,

Coastal currents in low latitudes

889

pressure gradient must balance the northward wind stress zY wherever the Coriolls force is too small to do so. In a shallow-water model, for example, the meridional momentum equation fu +

=

YlH

(1)

implies that the pressure gradient ~ly balances the wind stress at the equator y = 0, where the Coriolis parameter f vanishes. Far from the equator the meridional pressure gradient is unimportant and there is Ekman driR ~Y/fto the right of the wind. Figure 2 shows how this zonal Ekman drift changes near the equator. (The winds are assumed to be spatially uniform and zonal variations are neglected. Under such conditions there is no steady meridional flow in an inviscid shallow water model.) The zonal flow, which increases in amplitude as the equator is approached, forms a jet at a distance ~. from the equator: =

This is the equatorial radius of deformation. (The gravity wave speed is c = v / g H and 13is the latitudinal derivative of the Coriolis parameter. The coordinate ~ in Fig. 2 is a nondimensional latitudinal distance ~ = y/~..) After the sudden onset of winds it takes a time T, approximately, to reach the equilibrium state shown in Fig. 2. T = 1/X/pc

(3)

is the inertial time at a distance ~. from the equator. If the parameters are assigned reasonable oceanographic values then ~. is approximately 250 km and T is 2 days. (In models it takes about a week to reach equilibrium.) The principal difference between coastal zones near the equator and those in higher latitudes is the manner in which meridional (alongshore) gradients are established. Far from the equator coastal Kelvin waves introduce the gradients. Close to the equator offshore conditions impose meridional gradients on the coastal zone. Consider, for example, a northward coastal current driven by northward winds parallel to the western boundary of an ocean basin. In the southern hemisphere Ekman drift into the coast augments the transport of the current as it approaches the equator, but north of the equator the transport of the coastal current decreases, sharply so in the neighborhood of 3°N where the offshore Ekman drift is most intense. Could this be a factor in the separation of the Somali Current from the African coast near 4°N?

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890

5 G . H. PHILANDER a n d P. DEL.ECLUSE

This paper is a study of coastal zones in low latitudes. Specifically we describe the linear and nonlinear responses of a multi-level numerical model to the sudden onset and to the sudden relaxation of spatially uniform northward winds. THE

MODEL

The equations of motion (the p r i m i t i v e equations) are simplified by making the Boussinesq and hydrostatic approximations and by assuming an equation of state of the form p = po(1 -

aT),

where p is the density, T is the temperature, a = 0.0002(°C) -~ is the coefficient of thermal expansion, and P0 - 1 g cm -3. The coefficients of horizontal v x and vertical v V eddy viscosity and the thermal diffusivity r are assigned the values v r, = I0 c m 2 s -z

v . = 2 x 10 7 s -l

gv : l c m 2 S-t

K H--- 107 c m 2 S-I

except in the neighborhood of zonal boundaries, where the values of v u and Kn ate increased by a factor of 20 to dampen westward propagating Kelvin waves along those coasts. The model ocean is a rectangular box from I l ° S to 16°N with a longitudinal extent of 4600 kin. The latitudinal grid spacing is 55.6 kin; the longitudinal grid spacing is 111.2 km in the interior of the basin, but it decreases to 22.4 k m near the eastern and western coasts. (The grid spacing is evident in Figs 8, 11, and 12.) There are 16 grid points in the vertical. The ocean is 3000 m deep. Figure 3 shows the distribution of the upper grid points, and shows the initial temperature distribution. TEMPERATURE (°C) 5"

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Fig. 3. The initialtemperatureand Bruni-Viisilifrequency in the upper 500 m of the model. At greater depths the temperature decreases linearly to zero. The vertical spacing of grid ~ n t s is indicated.

Coastal currents in low latitudes

891

The ocean is at rest initially and is forced by an imposed wind stress of intensity 0.5 dyn cm -2 . The horizontal velocity components vanish at the vertical walls and at the ocean floor. The onset of the winds is gradual over a period of five d a y s - - a cosine taper is used--so as to filter out inertia-gravity waves. In the case where the winds relax after blowing steadily for 35 days this again happens over a period of five days. In linear calculations the equation for the conservation of heat is linearized with the temperature in Fig. 3 as the basic temperature field. The heat flux across all boundaries, including the surface, is zero. When the vertical temperature gradient becomes unstable, convective adjustment creates a mixed layer with zero vertical gradients. LINEAR RESPONSE

Figure 2 shows the response of a shallow-water model to northward winds when zonal variations are neglected. In the equilibrium state the zonal currents are westward south of the equator, eastward north of the equator, and there is no meridional flow'. The thermocline is depressed in the neighborhood of 3°N and is raised near 3°S. This implies a southward pressure force that balances the northwward wind stress near the equator. Consider next a stratified model. The wind stress now acts in a relatively shallow surface layer while the pressure force acts over a greater depth. The meridional circulation shown in Fig. 4 is therefore possible. The surface layers move northward in the direction of the wind; the subsurface layers, which include the thermocline, move southward. Downwelling near 3°N, where the thermocline is depressed, and upwelling near 3°S, where the thermocline is raised, close the circulation. The zonal currents at the surface are similar to those in Fig. 2: there is westward Ekman drift south of the equator; eastward drift north of it. Just below the surface, where the direct effect of the wind is secondary, the Coriolis force gives rise to weak equatorial currents at fight angles to the southward pressure force. The subsurface currents are westward north of the equator, eastward south of it (Fig. 4). (This paper describes currents driven directly by the wind or driven by pressure gradients maintained by the wind in the upper ocean. It does not concern deep currents below the thermocline.) In a shallow-water model equilibrium conditions evolve in the approximate time T given by equation (3). In a stratified ocean with Brunt-V/iisiil/i frequency N the gravity wave speed is

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Fig. 4. A meridional section of the two horizontal velocity components and temperature along a meridian in the center of the basin. The fields were averaged over the 13-day period from day 34 to day 46 to eliminate the Rossby-gravity oscillations. (The winds start to blow on day zero.) In shaded regions the flow is (a) westward or (b) southward.

s. G, 1t. PHILANDERand P. DELECLUSf-

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NH, where H is a vertical scale. Hence

T: i/v/ Nu. For the stratification being used in the model (Fig. 3), H is taken to be the depth of the thermocline and N is the value of the Brunt-Viiisiilii frequency in the thermocline. This gives a value for T of the order of two days. Figure 5 shows that equilibrium conditions are reached within a week. Oscillations with the period 2)iT = 13 days persist for a considerable time. These are Rossby-gravity waves with a zonal wavenumher equal to zero. They are excited by the spatially uniform winds. When the waves, which have eastward group velocities, impinge on the eastern coast they generate poleward propagating coastal Kelvin waves and thus lose energy. Consider next the evolution of equilibrium conditions near the coasts. Far from the equator the coastal jet initially accelerates until the arrival of Kelvin waves that establish pressure gradients that balance the wind stress. In low latitudes the coastaljets evolve as shown in Fig. 6. Along the western coast the acceleration of the jet ceases fn'st in the higher latitudes, consistent with the equatorward propagation of a Kelvin wave. Near the equator the Kelvin wave is difficult to discern because large latitudinal gradients are established independently of the Kelvin wave, and the wave has complex properties. (Near the equator the energy of coastal Kelvin waves is transferred to equatorially trapped waves.) Along the eastern coast a succession of poleward travelling Kelvin waves is evident in Fig. 6b. They are excited by the Rossby-gravity waves, with zero zonal wavenumher, that are incident on that coast. Because of these Kelvin waves the eastern coastal jet does not become steady the way the western jet does, but it oscillates about a mean value. Figure 7 shows that the coastal jet along the western boundary is far more intense than the jet along the eastern boundary. The primary reason for this is the difference in the meridional ~ pressure gradients on the two sides of the basin. In the middle of the basin the pressure force

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Fig. 7. Meridional sections of the alongshore velocity component and temperature 60 km from the western (a) and (c) and eastern (b) and (d) coasts. The fields were averaged over a 13-day period, from day 34 to day 46, to filter out the dominant oscillations. The flow is southward in shaded

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Coastal currents in low latitudes

895

at the equator is southward and balances the northward wind stress. Near the eastern coast the pressure force is augmented because the offshore depression of the thermocline near 3°N is enhanced at the coast--the winds cause downweUing north of the equator--and the offshore elevation of the thermocline near 3°S is also enhanced. This means that near the equator the northward surface current in Fig. 7b is flowing 'uphill' while the southward undercurrent is flowing 'downhill'. Along the western coast the situation is reversed; the winds cause coastal upwelling north of the equator, downwelling south of it. This can cause the pressure force along the western coast to be northward so that the surface current, when it crosses the equator, is flowing 'downhill.' Friction must therefore be invoked to achieve a momentum balance near the equator along the western coast. A further difference between the two coastal currents is the effect of the Ekman drift on them. Along the western coast the drift increases the transport of the current as it approaches the equator and decreases it north of the equator. The maximum transport is at the equator. The coastal jet in the east has a minimum at the equator because it loses fluid as it approaches the equator and gains fluid in the northern hemisphere. The final phase in the evolution of equilibrium conditions is associated with the westward propagation of long Rossby waves, at speeds c/5, c/9, and c/13 from the eastern coast. (The gravity wave speed c is 150 cm s-l.) The waves disperse features, initially confined to the neighborhood of the coast, westward. In an inviscid model the equilibrium state in the wake of all the waves is a state of no motion in which a southward pressure force balances the northward wind stress everywhere (CANE and SARACHIK, 1979). If dissipation is taken into account, the waves attenuate as they propagate westward. The currents described earlier therefore persist in the central and western part of the basin, as if no waves had been excited at the eastern coast. In the eastern part of the basin the waves modify the initial conditions. In Fig. 8, which shows various fields 50 days after the onset of southerly winds, a region that extends 15 ° longitude offshore from the eastern coast is affected by Rossby waves. For a detailed discussion of changes on longer time scales refer to PmLANDER and PACANOWSKI (1981). Rossby waves excited along the western coast are so slow, short, and prone to dissipation that they rarely cause any dispersion as is evident in Fig. 8. THE NONLINEAR

RESPONSE

A fluid parcel that is displaced poleward from the equator acquires eastward momentum provided it conserves its angular momentum. CANE (1979) invoked this argument to explain why nonlinearities intensify the eastward surface current north of the equator (Fig. 4) into a jet at a distance 2(xY/g3Zh)1/3 from the equator. (Here h is effectively the depth of the thermoeline.) For reasonable values of the parameters the distance is approximately 4 ° of latitude, essentially the same as the radius of deformation. It does, however, increase as the intensity of the wind stress increases. The meridional circulation is of central importance in this nonlinear mechanism so that it is absent from nonlinear shallow-water models but is present in the nonlinear constant-density model of CHARNEY and SPmGEL (1971). The latter model, which reproduces currents remarkably similar to those in Fig. 9, has vertical friction so that a meridional circulation is possible. (The wind acts as a surface stress, not as a body force.) Figure 9 shows that because of advection, the westward surface current, which is south of the equator in linear models, penetrates into the northern hemisphere where it merges with the subsurface westward current (Fig. 4) and hence appears to penetrate below the thermodine.

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Along the western boundary, nonl~arities cause the maximum of the coastal jet to move from the equator into the northern hemisphere (Fig. 10). UpwcUing is most intense just north of the region of maximum northward velocities (Figs 10 and 11). At subsurface levels, a clockwise gyre forms off the western coast in the northern hemisphere. Along the eastern side of the basin the effect of nonlinearities is less pronounced than along the western side.

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linear model the eastward flow in the northern hemisphere and the westward flow in the southern hemisphere (Fig. 8) disappear within a week or two after the winds stop blowing. In a nonlinear model, however, the zonal currents in Figs 7 and 11 adjust towards a geostrophic balance with the density field. In the absence of friction such currents persist until Rossby waves from the eastern coast eliminate them. Along the western boundary there is a region where the pressure force is in the direction of the coastal jet. In this region a relaxation of the wind barely affects the coastal jet, which is gradually decelerated by frictional forces (Fig. 10). In the coastal zone just south of the equator, the pressure force opposes the northward jet and causes it to reverse direction when the wind stops blowing. The development of such a southward current when the northward winds relax is clearly evident in Fig. 10. CONCLUSIONS AND DISCUSSION

The results suggest the following response to the sudden onset of the southwest monsoons over the Indian Ocean in May. An accelerating jet appears immediately all along the coast of east Africa. Poleward of 5 o latitude the acceleration stops in the wake of equatorward travelling coastal Kelvin waves that establish alongshore pressure gradients to balance the wind. In the equatorial zone, alongshore pressure gradients are primarily due to the strong windinduced upwelling near 3°N, where the offshore Ekman drift is most intense, and the strong downwelling near 3 °S, where there is large onshore Ekman drift. The deformation of the thermocline results in a pressure force in the direction of the wind so that the coastal current accelerates to a high speed and friction must be invoked to establish equilibrium conditions. [Cox (1976, 1979) found the intensity of the Somali Current to be very sensitive to the parameterization of friction in his model.] Equilibrium conditions obtain approximately two weeks after the sudden onset of the winds. This is the time it takes Kelvin waves to reach the equator from the northern and southern boundaries of the basin, and it is also the local frictional time scale, (8/132A)~3, where A is the coefficient of horizontal eddy viscosity (HuRLBURT and THOMPSON, 1976). By the time the coastal jet is steady it is highly nonlinear so that advection has moved the core of the jet from the equator to approximately 4°N, where the jet veers offshore. This is the latitude where nonlinearities intensify the offshore Ekman drift into an eastward jet. Near the equator, the Somali Current can therefore be viewed as a western boundary current that serves to transport the Ekman drift into the coast south of the equator, into the northern hemisphere where it is lost to the eastward current near 4°N. Upwelling is intense in the separation region so that a tongue of cold surface waters extends offshore in this region. The feature is evident in Fig. 11. If the eastward component of the southwest monsoon and the inclination of the African coast to meridians are taken into account, then the tendency of the coastal jet to overshoot its point of separation is inhibited. Simulations of the Somali Current with nonlinear numerical models that include these factors are very realistic (DELECLUSE,1983). According to the argument advanced above, the coastal jet separates from the coast because of constraints imposed by offshore conditions. Specifically, the coastal jet veers eastward near 3°N because it must feed the eastward current maintained by the crossequatorial winds. There is of course a vorticity balance at all points in the basin, in particular in the region where the jet separates from the coast. But this does not mean that the separation is attributable to constraints imposed by the vorticity balance. (Similarly, the separation of the Gulf Stream from the coast is dictated by the large-scale wind field--separation occurs where the curl is zero--not by local vorticity considerations.)

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In September and October the southwest monsoons weaken and die out. Our calculations suggest that when this happens, the current reverses direction south of the equator because in that region the pressure force that opposes the wind is a source of southward momentum. Further north where the jet separates from the coast, the pressure force is in the direction of the wind. Hence when the wind weakens, the current decelerates very gradually, primarily because of frictional effects. The tongue of cold water does not drift northward in the model when the winds relax so that there must be another explanation for the observed phenomenon. The calculations described here are for spatially uniform winds, but such winds are unrealistic. In reality the intensity of the coastal winds increases in a northward direction and has a maximum value near IO°N. Such spatial variations, which need not be associated with a curl, could drive an anticyclonic gyre along the northern part of the east African coast (D. L. T. ANDERSON,personal communication). In the absence of any curl the winds parallel to the coast could therefore drive both the southern and northern branches of the Somali Current. The measurements of SCHOTTand QUAOFASEL(1983) however show that the curl of the wind just offshore is also a driving force for the northern gyre. For an indication of the effects of the large-scale wind stress curl we used the observed mean monthly winds (HELLEgMAN, 1983) and SwgnRtJp's (1947) model to calculate the streamlines for the vertically integrated transports. The results, shown in Fig. 13 for May, June, and July, indicate that the coastal current does not separate from the coast near 4 o N. In other words, the curl of the wind is not responsible for that feature of the observed flow. Instead, the curl drives a coastal current that is continuous along the coast and which, according to Fig. 13, is most intense in July. There will be a delay of a few weeks at least before the circulation in Fig. 13 is established. Hence the effect of the curl of the wind on the coastal current will be most pronounced in August and September. The curl could therefore be responsible for the observed northward drift of the cold tongue, which is initially at 4°N. Cox (1976) described the response of a model in which the curl of the wind is associated with an abrupt decay of the northward wind stress along a line 1000 km east of the western boundary. The curl forces Somali Current 'eddies' northward. When the curl is associated with the abrupt decay of southwesterly winds along a line 1000 km offshore from a coast inclined at 45 ° to meridians, then the separation region does not drift northward. This suggests that migrations of the tongue of cold water near 4 ° N depend on the details of the imposed curl. Given the considerable interannual variability of the surface winds in the Indian Ocean, the tongue could behave very differently in different years. The explanations given here for various features of the Somali Current depend on results obtained with an idealized model. The stage now seems set for a more ambitious study that takes into account that the ocean is not at rest when the southwest monsoon starts to blow--ANDERSON and MooRE (1979) have proposed that the southern hemisphere winds drive a Somali Current before the local winds do---and that takes into account the complex spatial and temporal variations of the winds. The winds off the western coasts of South America and Africa vary primarily on long time scales and hence drive the broad slow Peruvian and Benguela currents (PalLANOER and PACANOWSgl, 1981). The southeast trades do however have some short-term variability. Of particular interest is the seasonal relaxation of such winds in February and March. If this happens over a short period, then the southward pressure force, which the winds had maintained, is left unbalanced and a warm southward coastal current results. We propose this as an explanation for the annual El Nifio current along the coasts of Ecuador and northern Peru. The coastal current can persist for at most a month or two before Rossby waves cause

Coastal currents in low latitudes

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Fig. 13. The stream function for the vertically integrated horizontal transport, obtained by integrating Sverdrup's equation 13¥x= curlzT subject to the condition that ~ vanishes along the eastern coasts of the Indian Ocean. (It is assumed that there is land between Africa and Madagascar.)

westward dispersion, but by that time the wind would have intensified again, amplifying the southward pressure force, which then drives a subsurface southward current. Variability along the eastern coast is likely to be high, especially on time scales from a week to months. In this frequency range, which excludes inertia-gravity and Rossby waves, the only equatorial waves that are possible--Kelvin and Rossby-gravity w a v e s - - h a v e eastward group velocities and hence are incident on the eastern coast, where they excite polewardtravelling coastal Kelvin waves. (It is for the same reason that variability along the western boundary is low at periods from a week to a few months.) The calculations described here demonstrate how effectively cross-equatorial winds can excite coastal Kelvin waves along an eastern boundary. Fluctuations o f the zonal component of the wind will add a further degree of complexity, particularly so in February and March when a weakening of the westward winds causes the Equatorial Undercurrent to surge into the region east o f the Galapagos

902

s G. t|. PtlILANDER and P. DELECLUSE

I s l a n d s (LuKAS, 1982). M e a s u r e m e n t s n o w b e i n g m a d e in the e a s t e r n t r o p i c a l Pacitic O c e a n as p a r t o f a p r o g r a m called E P O C S ( E q u a t o r i a l Pacific O c e a n C l i m a t e Studies) s h o u l d a d d s u b s t a n t i a l l y to o u r k n o w l e d g e o f v a r i a b i l i t y in t h e e a s t e r n Pacific.

Acknowledgements--One of us (P.D.) was supported through Geophysical Fluid Dynamics Laboratory/NOAA Grant 04-7-022-44017, We are indebted to Ms. WILLIAMSfor typing the manuscript and to Mr. TUNISON and his staff for drafting the figures. REFERENCES ALLEN J. S. (1976) Some aspects of the forced wave response of stratified coastal regions. Journal of Physical Oceanography, 6, 113-119. ANDERSON D. L. T. and D. W. MOORE 0979) Cross-equatorial inertial jets with special relevance to very remote forcing of the Somali Current. Deep-Sea Research, 26, i-22. BROWN D. B., J. G. BRUCE and R. H. EVANS 0980) Evolution of sea surface temperature in the Somali Basin during the southwest monsoon of 1979. Science, Wash., 209, 595-597. CANE M. A. (1979) The response of an equatorial ocean to simple wind-stress patterns: lI. Numerical results. Journal of Marine Research, 37, 253-299. CANE M. A. and E. S. SARACHIK 0979) Forced baroclinic ocean motions Ill. Journal of Marine Research, 37. 355-398. CHARNEY J. G. (1955) The generation of oceanic currents by winds. Journal of Marine Research, 14, 477-498. CHARNEY J. G. and S. L. SPIEGEL (1971) Structure of wind.driven equatorial currents in homogeneous oceans. Journal of Physical Oceanography, l, 149-160. Cox M. D. (1976) Equatorially trapped waves and the generation of the Somali Current. Deep-Sea Research, 23, 1139-1152. Cox M. D. (1979) A numerical study of Somali Current eddies. Journal of Physical Oceanography, 9, 311-326. DELECLUSE P. (1983) On the dynamics of western boundary currents in low latitudes. (Submitted to Journal of

Physical Oceanography.) DOING W., R. L. MOLINARI and J. C. SWALLOW(1980) Somali Current: evolution of the surface flow. Science~ Wash., 209, 588-590. HELLERMAN S. (1983) Normal monthly windstress over the world ocean with error estimates. (Submitted to

Journal of Physical Oceanography.) HURLBURT H. E. and J. D. THOMPSON 0976) A numerical model of the Somali Current. Journal of Physical Oceanography, 6, 646-664. LUKAS R. (1982) The termination of the Equatorial Undercurrent in the eastern Pacific Ocean. Ph.D. thesis, University of Hawaii, 127 pp. MOORE D. W. and S. G. H. PHILANDER (1977) Modeling of the tropical oceanic circulation. In: The sea, Vol. 6, Wiley lnlerscience, New York, pp. 3t9-361. PHILANDER S. G. H. and R. C. PACANOWSK]( 1981) The oceanic response to cross-equatorial winds (with application to coastal upwelling in low latitudes). Tellus, 33, 201-210. SCHOTT F. and D. R. QUADFASEL(1983) LOw frequency current fluctuations off Somalia during the 1979 summer monsoons. Journal of Physical Oceanography(in press). SVERDRUP H. U. (1947) Wind-driven currents in a baroclinic ocean: with application to the equatorial currents in the eastern Pacific. Proceedings of the National Academy of Science, 33, 318-326. SVERDRUP H. U., M. W. JOHNSON and R. H. FLEMING (1942) The oceans. Prentice-Hall, Englewood Cliffs, New Jersey, 1087 pp.