Lectures - Monday and Wednesday, 11:00 AM - 12:15 PM
Lab - Tuesday, 4:10 PM -7 PM
Lecture Figures (powerpoint)
The ocean would have a significant role in governing climate, even if it did not circulate. The ocean’s surface layer ability to store heat in summer and the release of that heat to the atmosphere in winter, would mitigate the seasonal extremes of the atmosphere temperature even without an ocean circulation. The circulation shapes the sea surface temperature (SST) pattern moving heat from the tropics to the polar regions (Fig. 1). Currents along the western boundary of the ocean, such as the Gulf Stream are swift and form filaments of warm water projecting towards higher latitude. Ocean currents also affect the salinity patterns, as the ocean circulation moves saline water formed in the excess evaporative regions to the excess precipitation regions, subtropical to tropical respectively, as part of the global hydrological cycle.
Ocean circulation includes currents, which follow along an approximate horizontal plane, such as flow along the sea surface, and flow in the vertical. Ocean currents may be slow, 0.01 m/s in the ocean interior or swift >1.0 m/s as in the Gulf Stream. Movement of water in the vertical is much smaller, ranging from a maximum of 0.001 m/s to more characteristic speeds of 0.0001 m/s. The ocean circulation, both horizontal and vertical, is induced by two means (Fig. 2): (1) by the wind exerting a stress upon the sea surface, and (2) by buoyancy fluxes between the ocean and atmosphere. The former is called the wind driven circulation, the latter is the thermohaline circulation. There are also tidal currents, products of the Earth/Moon/Sun gravitational interaction. While we do not include discussion of the ocean tidal currents in this course, some oceanographers consider the dissipation of tidal energy within the ocean as an important factor in mixing the ocean and in shaping the SST patterns.
The wind driven circulation is by far is the more energetic, though confined mostly to the upper kilometer or two of the ocean. The mean surface winds patterns (Fig. 3) are composed of specific elements. The trade winds blow westward with a component of flow towards the equator. The trade winds of the two hemispheres meet at the intertropical convergence zone (ITCZ), where updrafts of air induce high precipitation. The ITCZ falls over the warmest band of SST, which on annual average occurs in the northern hemisphere, near 5°N. Poleward of the trades are the westerlies (air flow towards the east) with a slight poleward component, and still further poleward are the polar easterlies. The westerlies are associated with the strongest wind, but often very variable, as shaped by storms and atmospheric fronts.
The wind exerts a force or stress on the ocean surface, proportional to the square of the wind speed. This produces not just ocean waves but also injects momentum into the surface layer of the ocean. The wind makes the surface layer of the ocean move, though not in the way that intuition might dictate - it’s not in the direction of the wind stress, but rather at an angle to it, in a “spiral” across the surface layer (Fig. 4). This is because of the Coriolis Force. A balance is achieved between the wind stress and the Coriolis Force. The surface Ekman layer (named after the person who developed the theory in 1908) extends to about 50 to 200 meters depth. The mean transport within the Ekman layer is 90° towards the right of the wind in the northern hemisphere, 90° to the left in the southern hemisphere. Typically the wind induced surface current is around 2 or 3% of the wind speed. One clear effect of the Ekman transport can be seen in along the coast of California (Fig. 5), where cold subsurface water is brought up to the sea surface from a depth of perhaps 100 meters to replace surface water forced offshore by the Ekman transport. These regions are rich in nutrients and support important fisheries. Ekman divergence occurs along the Equator (Fig. 6) and Ekman convergence occurs in the subtropics (Fig. 7).
As we know the wind changes in strength and direction from place to place. This causes a spatially variable Ekman transport. In some regions the Ekman transport forces accumulation (convergence) of surface water, in other regions it results in the remove (divergence) of surface water. As surface water is less dense than deeper water (the ocean is stable, less dense more buoyant water at top, denser water at the bottom) this has the effect of heaping buoyant surface water in the convergence regions and removing it from the divergence regions. The buoyant surface layer thickens in convergence, thins in divergences regions. Through the Ekman transport mechanism the wind redistributed the buoyant surface layer of the ocean.
The thickening or thinning of the buoyant surface layer by the Ekman transport produces “hill” and “valleys” of the sea surface (Fig. 8). The total relief of the ocean surface (not counting waves or tides) is only about 1.5 meters, but this is sufficient to drive the ocean circulation. To understand how the wind stress distorts sea level, with hills at the convergences, valleys at the divergences, one needs to invoke the hydrostatic relationship. Hills of buoyant water are balanced by a 'root' of buoyant water (Fig. 9), much as with a ship or iceberg floats, with the part above the sea level is balanced by the submerged part. Removal of buoyant water in the Ekman divergences induces a depression in sea level and upward bowing of the deeper dense water. The vertical movement of water at the base of the Ekman Layer act to distort the density fields within the deeper water.
How do hills and valley drive a circulation? The pattern of the induced currents is governed not just by the wind stress and its Ekman transport but also by the earth's rotation, the Coriolis Force. The Coriolis Force acts at right angles (90°) to the ocean current (or wind) direction, to the right in the north hemisphere to the left in the southern hemisphere. The Coriolis Force strength is proportional to the strength of the ocean current. In short: the wind produces convergences and divergences of surface water, which causes hills and valleys of sea level, which then produce a horizontal gradient of pressure. Hills have an outward directed pressure gradient, depression an inward directed pressure gradient (Fig. 10). As the pressure gradients make the water move from high pressure to low pressure, the Coriolis Force starts its action, and eventually a balance is achieved in these two forces, the horizontal pressure gradient equals the magnitude of the Coriolis Force, but is directed in the opposite direction. This balance is called the geostrophic balance, and a current in such a balance is called a geostrophic current. Ocean currents are very close to being in geostrophic balance (Figs. 11-13).
The exchange of heat and freshwater between the ocean and atmosphere alters the density of the surface water. Cooling of the ocean makes the ocean denser, removing buoyancy; evaporation makes the ocean saltier and hence denser, again removing buoyancy. Heating and excess precipitation has the opposite effect, they add buoyancy to the ocean. Surface water that is made denser, through cooling or increase of salinity, sinks to an equilibrium depth (a depth where its density matches that of the ambient ocean). It then spreads into the world ocean, displacing the older resident water made less dense by ocean mixing. To reach a steady state condition the loss of water from the surface layer must be replaced, the loop must be closed: the displaced water must find its way back to the sinking region. This sets up the thermohaline circulation. The thermohaline circulation is more sluggish than the wind driven circulation, generally less than 0.1 m/s. But it involved overturning of the full depth of the ocean, hence bringing deep water removed from direct contact with the atmosphere for 100 to 1000 years back to sea surface.
A key part of the thermohaline circulation is the production of North Atlantic Deep Water. Warm/salty surface water spread into the northern north Atlantic (Figs. 14, 15), where cooling occurs. The dense water sinks, overflows the ridges spanning the distance from Greenland to Iceland (Fig. 16) to spread as a salty water mass into the world ocean (Figs. 17, 18).
Along the margins of Antarctica very dense water is formed (Figs 19, 20). This water descends as thin sheets or plumes down the continental slope into the deep ocean forming Antarctic Bottom water that spreads along the sea floor over much of the world ocean.
Text by Arnold Gordon, 2004.