14.1: Equatorial Processes

The tropical ocean is characterized by a thin, permanent, shallow layer of warm water over deeper, colder water. In this respect, the vertical stratification is similar to the summer stratification at higher latitudes. Surface waters are hottest in the west (figure \(6.3.2\)) in the great Pacific warm pool. The mixed layer is deep in the west and very shallow in the east (Figure \(\PageIndex{1}\)).

Figure \(\PageIndex{1}\): The mean upper-ocean thermal structure along the equator, in the Pacific from north of New Guinea to Ecuador. Calculated from data in Levitus (1982).

The shallow thermocline has important consequences. The southeast trade winds blow along the equator (figure \(4.2.1\)), although they tend to be strongest in the east. North of the equator, Ekman transport is northward. South of the equator it is southward. The divergence of the Ekman flow causes upwelling on the equator. In the west, the upwelled water is warm, but in the east the upwelled water is cold because the thermocline is so shallow. This leads to a cold tongue of water at the sea surface extending from South America to near the dateline (figure \(6.3.2\)).

Surface temperature in the east is a balance among four processes:

1. The strength of the upwelling, which is determined by the westward component of the wind.
2. The speed of westward currents which carry cold water from the coast of Peru and Ecuador.
3. North-south mixing with warmer waters on either side of the equator.
4. Heat fluxes through the sea surface along the equator.

The east-west temperature gradient on the equator drives a zonal circulation in the atmosphere, the Walker circulation. Thunderstorms over the warm pool carry air upward, and sinking air in the east feeds the return flow at the surface. Variations in the temperature gradient influences the Walker circulation, which, in turn, influences the gradient. The feedback can lead to an instability, the El Niño-Southern Oscillation (ENSO) discussed in the next section.

Surface Currents

The strong stratification confines the wind-driven circulation to the mixed layer and upper thermocline. Sverdrup’s theory and Munk’s extension, described in Section 11.1 and Section 11.3, explain the surface currents in the tropical Atlantic, Pacific, and Indian ocean. The currents include (figure \(\PageIndex{2}\)):

1. The North Equatorial Countercurrent between 3\(^{\circ}\)N and 10\(^{\circ}\)N, which flows eastward with a typical surface speed of 50 cm/s. The current is centered on the band of weak winds, the doldrums, around 5–10\(^{\circ}\)N where the north and south trade winds converge, the tropical convergence zone.
2. The North and South Equatorial Currents, which flow westward in the zonal band on either side of the countercurrent. The currents are shallow, less than 200 m deep. The northern current is weak, with a speed less than roughly 20 cm/s. The southern current has a maximum speed of around 100 cm/s, in the band between 3\(^{\circ}\)N and the equator.

Figure \(\PageIndex{2}\): Average currents at 10 m calculated from the Modular Ocean Model driven by observed winds and mean heat fluxes from 1981 to 1994. . The model, operated by the NOAA National Centers for Environmental Prediction, assimilates observed surface and subsurface temperatures. After Behringer, Ji, and Leetmaa (1998).

The currents in the Atlantic are similar to those in the Pacific because the trade winds in that ocean also converge near 5–10\(^{\circ}\)N. The South Equatorial Current in the Atlantic continues northwest along the coast of Brazil, where it is known as the North Brazil Current. In the Indian Ocean, the doldrums occur in the southern hemisphere and only during the northern-hemisphere winter. In the northern hemisphere, the currents reverse with the monsoon winds.

There is, however, much more to the story of equatorial currents.

Equatorial Undercurrent: Observations

Just a few meters below the surface on the equator is a strong eastward flowing current, the Equatorial Undercurrent, the last major oceanic current to be discovered. Here’s the story:

In September 1951, aboard the U.S. Fish and Wildlife Service research vessel long-line fishing on the equator south of Hawaii, it was noticed that the subsurface gear drifted steadily to the east. The next year Cromwell, in company with Montgomery and Stroup, led an expedition to investigate the vertical distribution of horizontal velocity at the equator. Using floating drogues at the surface and at various depths, they were able to establish the presence, near the equator in the central Pacific, of a strong, narrow eastward current in the lower part of the surface layer and the upper part of the thermocline (Cromwell, et. al., 1954). A few years later the Scripps Eastropac Expedition, under Cromwell’s leadership, found the current extended toward the east nearly to the Galapagos Islands but was not present between those islands and the South American continent.

The current is remarkable in that, even though comparable in transport to the Florida Current, its presence was unsuspected ten years ago. Even now, neither the source nor the ultimate fate of its waters has been established. No theory of oceanic circulation predicted its existence, and only now are such theories being modified to account for the important features of its flow.—Warren S. Wooster (1960).

The Equatorial Undercurrent in the Atlantic was first discovered by Buchanan in 1886, and in the Pacific by the Japanese Navy in the 1920s and 1930s (McPhaden, 1986).

However, no attention was paid to these observations. Other earlier hints regarding this undercurrent were mentioned by Matthäus (1969). Thus the old experience becomes even more obvious which says that discoveries not attracting the attention of contemporaries simply do not exist.— Dietrich et al. (1980).

Bob Arthur (1960) summarized the major aspects of the flow:

1. Surface flow may be directed westward at speeds of 25–75 cm/s;
2. Current reverses at a depth of from 20 to 40 m;
3. Eastward undercurrent extends to a depth of 400 meters with a transport of as much as \(30 \ \text{Sv} = 30 \time 10^{6} \ \text{m}^{3}/\text{s}\);
4. Core of maximum eastward velocity (0.50–1.50 m/s) rises from a depth of 100 m at 140\(^{\circ}\)W to 40 m at 98\(^{\circ}\)W, then dips down;
5. Undercurrent appears to be symmetrical about the equator and becomes much thinner and weaker at 2\(^{\circ}\)N and 2\(^{\circ}\)S.

In essence, the Pacific Equatorial Undercurrent is a ribbon with dimensions of \(0.2 \ \text{km} \times 300 \ \text{km} \times 13,000 \ \text{km}\) (figure \(\PageIndex{3}\)).

Figure \(\PageIndex{3}\): Cross section of the Equatorial Undercurrent in the Pacific calculated from Modular Ocean Model with assimilated surface data (See Section 14.5). The section is an average from 160\(^{\circ}\)E to 170\(^{\circ}\)E from January 1965 to December 1999. Stippled areas are westward flowing. From Nevin S. Fučkar.

Equatorial Undercurrent: Theory

Although we do not yet have a complete theory for the undercurrent, we do have a clear understanding of some of the more important processes at work in the equatorial regions. Pedlosky(1996), in his excellent chapter on Equatorial Dynamics of the Thermocline: The Equatorial Undercurrent, points out that the basic dynamical balances we have used in mid-latitudes break down near or on the equator.

Near the equator:

1. The Coriolis parameter becomes very small, going to zero at the equator: \[f = 2 \Omega \sin \varphi = \beta y \approx 2 \Omega \varphi \nonumber \]where \(\varphi\) is latitude, \(\beta = \partial f/\partial y \approx 2 \Omega/R\) near the equator, and \(y = R \varphi\).
2. Planetary vorticity \(f\) is also small, and the advection of relative vorticity cannot be neglected. Thus the Sverdrup balance \((11.1.10)\) must be modified.
3. The geostrophic and vorticity balances fail when the meridional distance \(L\) to the equator is \(O \left(\sqrt{U/\beta}\right)\), where \(\beta = \partial f/\partial y\). If \(U = 1 \ \text{m/s}\), then \(L = 200 \ \text{km}\) or 2\(^{\circ}\) of latitude. Lagerloeff et al (1999), using measured currents, show that currents near the equator can be described by the geostrophic balance for \(|\varphi| > 2.2^{\circ}\). They also show that flow closer to the equator can be described using a \(\beta\)-plane approximation \(f = \beta y\).
4. The geostrophic balance for zonal currents works so well near the equator because \(f\) and \(\partial \zeta/\partial y \rightarrow 0\) as \(\varphi \rightarrow 0\), where \(\zeta\) is sea surface topography.

Upwelled water along the equator produced by Ekman pumping is not part of a two-dimensional flow in a north-south, meridional plane. The flow is three-dimensional. Water tends to flow along the contours of constant density (isopycnal surfaces), close to the lines of constant temperature in figure \(\PageIndex{1}\). Cold water enters the undercurrent in the far west Pacific, and it moves eastward and upward along the equator. For example, the 25\(^{\circ}\) isotherm enters the undercurrent at a depth near 125 m in the western Pacific at 170\(^{\circ}\)E and eventually reaches the surface at 125\(^{\circ}\)W in the eastern Pacific.

The meridional geostrophic balance near the equator gives the speed of the zonal currents, but it does not explain what drives the undercurrent. A very simplified theory for the undercurrent is based on a balance of zonal pressure gradients along the equator. Wind stress pushes water westward, producing the deep thermocline and warm pool in the west. The deepening of the thermocline causes the sea-surface topography \(\zeta\) to be higher in the west, assuming that flow below the thermocline is weak. Thus there is an eastward pressure gradient along the equator in the surface layers to a depth of a few hundred meters. The eastward pressure gradient at the surface (layer A in figure \(\PageIndex{4}\)) is balanced by the wind stress \(T_{x}\), and \(T_{x}/H = -\partial p/\partial x\), where \(H\) is the mixed-layer depth.

Figure \(\PageIndex{4}\): Left: Cross-sectional sketch of the thermocline and sea-surface topography along the equator. Right: Eastward pressure gradient in the central Pacific caused by the density structure at left.

Below a few tens of meters in layer B, the influence of the wind stress is small, and the pressure gradient is unbalanced, leading to an accelerated flow toward the east, the equatorial undercurrent. Within this layer, the flow accelerates until the pressure gradient is balanced by frictional forces which tend to slow the current. At depths below a few hundred meters in layer C, the eastward pressure gradient is too weak to produce a current, so \(\partial p/\partial x \approx 0\). Coriolis forces keep the equatorial undercurrent centered on the equator. If the flow strays northward, the Coriolis force deflects the current southward. The opposite occurs if the flow strays southward.