Figure 7. The northward transport of energy by ocean and atmosphere as a function of latitude (averaged at each latitude around the globe over a year). Negative values denote southward transport.
The ocean is as important as the atmosphere in transporting heat. Also the effectiveness of doing so strongly depends on the specific mechanisms. There is a large difference between atmosphere and ocean itself. Ocean currents that carry warm water from the tropics towards higher latitudes are very efficient in transporting heat at latitudes of about 20° (Figure 7).
Figure 7. The northward transport of energy by ocean and atmosphere as a function of latitude (averaged at each latitude around the globe over a year). Negative values denote southward transport.
Displaced water warms the air and, indirectly, the land over which the air blows. As such land areas near oceans have maritime climates, which are less extreme than continental ones, with smaller day-night and winter-summer differences. The surface currents in the ocean are forced by atmospheric surface winds. Although there is a large similarity in patterns, the connection is not as obvious as it might first appear. One difference is due to the shapes of the ocean basins. Hence, the tendency for circular (gyral) motion is even more noticeable in the oceans, e.g., the North Pacific gyre, which includes the Kuroshio, the North Pacific current, the California current and the North Equatorial current (Figure 8).
Another outstanding difference is the complex current system near the equator. The equatorial counter-currents, i.e., currents against the prevailing wind direction, are small and sometimes not very good developed. Note that Figure 8 represents average conditions and that differences between ocean currents and wind fields are most evident during different seasons. The oceanic counterparts of transient weather systems (e.g., cyclones) are called mesoscale eddies. They last much longer than atmospheric (anti)cyclones and are also smaller.
So far, we have looked at the surface of the ocean. However, vertical circulation cells, like the Hadley cell in the atmosphere, are also possible in the ocean. Vertical motion in the atmosphere is related to density differences. However, these differences are only effective in producing motion because of gravity. In reality differences in buoyancy, i.e., the weight per unit volume, are important. A particle is said to be more buoyant when it has less weight. The ocean also moves because of buoyancy contrasts, but these differences are due to salinity (i.e., concentration of dissolved salt) as well as to temperature. The ocean circulation that is driven by buoyancy differences is called the thermohaline circulation.
Salinity is expressed in ‰ or concentration in parts per thousand by weight. The major dissolved constituents are chloride, sodium, sulphate and magnesium ions, respectively. Salinity in surface waters of the open ocean ranges from 33 to 37. The way salinity varies throughout the oceans depends almost entirely on the balance between evaporation and precipitation (Figure 9). The salinity of the surface waters is at a maximum in latitudes of about 20°, where evaporation largely exceeds precipitation. Remember that these regions are close to the descending part of the Hadley cell. Salinities decrease both towards higher latitudes and towards the equator.
Figure 8 The annual mean surface currents in the oceans. There are seasonal differences in most areas, particularly in areas that experience strong seasonal variations in wind directions. Currents on the Western side of the ocean are 50-100km wide, the Eastern currents are even wider but also slower.
Figure 9 Average values of salinity and the differences between annual mean evaporation and precipitation (E-P) as function of latitude.
The way both temperature and salinity influence buoyancy is very important to the thermohaline circulation. Figure 10 shows that the freezing point and the temperature of maximum density are the same when salinity of water reaches about 25. Hence, the density of seawater increases with increasing salinity and falling temperature right down to the freezing point. Evaporation at the ocean surface, for example, decreases buoyancy by cooling and by increasing salinity. Precipitation, meltwater and rivers increase buoyancy locally by decreasing salinity.
Figure 10 Temperature of freezing point and maximum density of water as functions of salinity.
Also the formation of sea-ice influences buoyancy locally. When sea-ice is formed, almost all the salt is left in the remaining seawater, which thus becomes more saline. This cold seawater with higher salinity is very dense, and sinks to greater depths. In this way some of the densest water is formed near Antarctica. This Antarctic Bottom Water flows along the bottom of the ocean around Antarctica and out into the Atlantic, Pacific and Indian Oceans (Figure 11). Near the North Pole, North Atlantic Deep Water (NADW) is formed mainly between Greenland and Norway, and flows south along the bottom of the Atlantic Ocean. NADW flows over Antarctic Bottom Water because it is less dense (Figure 11).
Figure 11 Meridional cross-section of the Atlantic Ocean, showing temperature- and salinity distribution. Bodies of water are identifiable by, among others, temperature and salinity for the very reason that mixing in the deep ocean occurs very slowly. Water masses that form in semi-enclosed seas provide particularly clear examples of bodies of water with recognisable temperature and salinity characteristics. Mediterranean Water (Medit.) is distinguished from other water masses by its relative high temperature and high salinity, and therefore can be recognised throughout much of the Atlantic Ocean at a depth of about 1000m, where it is neutrally buoyant. Intermediate water (e.g., Antarctic Intermediate Water; Ant. int.) is formed in subpolar regions where precipitation exceeds evaporation, and its salinity is therefore low. NADW and Ant. bot. refer to North Atlantic Deep Water and Antarctic Bottom Water, respectively.
Measurements of the characteristics of water masses reveal a gigantic ocean 'conveyor belt' of deep water driven by the dense water sinking in the North Atlantic region (Figure 12). On timescales of about a millennium this conveyor includes the transport of deep water from the North and South Atlantic into the Indian and Pacific Oceans, and the return flow of relative warm water near the surface (Figure 12). This part of the thermohaline circulation transfers heat from the North Pacific and Indian Ocean into the North Atlantic. As such it has large influences on local climate over much of the Northern Hemisphere. Note that the thermohaline and ocean surface circulation are interconnected. When the NADW spreads to the bottom of the sea, and begins to flow towards the equator, this must necessarily intensify the surface current in the opposite direction. Hence, NADW warms Northwest Europe in two ways, first by transporting 'cold' to the deeper ocean, second by intensifying warm water transport of the North Atlantic drift.
Figure 12 The long term thermohaline circulation or 'conveyor belt'. NADW transports cold water towards the North Pacific and Indian Oceans. The narrow Agulhas Current near South Africa (Figure 2.9) is critical for the return flow of heat through surface currents from the Indian Oceans to the Atlantic Ocean. The heat and fresh water fluxes between these two oceans take place largely by Agulhas rings that pinch-off from the Agulhas Current and penetrate the Atlantic. This transport of heat associated with this conveyor belt fluctuates substantially over time-scales ranging from years to millennia.
Links
De Indisch-Atlantische Oceaan Connectie, Agulhasstroom