Winddriven circulation of the ocean

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In the last three sections we have seen how the wind acts directly on the ocean and the local effect of ocean surface drag on the atmosphere. The momentum transferred to the water, plus the heat and fresh water, creates pressure gradients in the ocean leading to motion on a larger scale - the thermohaline circulation. This will be considered in §2.12. The larger scale impact of the ocean on the atmosphere is also through the creation of pressure gradients, but via surface heating or cooling. The Coriolis force then acts as well, to give geostrophic balance and maintain the general circulation in each fluid. In this section we will discuss these basin-scale flows in the ocean while in §2.13 and Chapter 5 we will discuss the large-scale impact of surface heating on the marine atmosphere.

2.11.1 The ocean gyres

The key to understanding ocean gyres is contained in the concept of angular momentum that we discussed in §2.6.2. Conservation of angular momentum led to the Coriolis force. The winds of the sub-tropical anticyclones provide the ocean with angular momentum. As these winds blow semi-permanently the oceanic sub-tropical gyres should gradually accelerate. They do not; the gyre circulation is stable. Therefore there must be a mechanism providing a source of oppositely rotating angular momentum to balance the wind's influence.

Meteorologists and oceanographers use the concept of vorticity to replace angular momentum because there are two ways a body of water can gain angular momentum, and vorticity combines them. As we saw in our discussion of the Coriolis force, a body of water will change its angular momentum relative to the Earth's axis when moving north or south. This type of angular momentum is called planetary vorticity. We have also seen several situations where water rotates about a local axis of rotation, for example in inertial oscillations or an ocean gyre. The angular momentum possessed due to this local rotational motion is relative vorticity. This difference is illustrated schematically in Fig. 2.31.

A body of water has depth as well as breadth. Thus if the rotation rate increases, the net effect of the decrease in radius needed to conserve relative vorticity is to increase the depth (so as to conserve mass), as shown in Fig. 2.32. Conservation of net angular momentum, or vorticity, is therefore an interplay of several effects, described by conservation of the quantity (f + Z)/D, where

Fig. 2.31. Schematic showing the two forms of vorticity for a location in the mid-latitudes of the Northern Hemisphere.

Fig. 2.31. Schematic showing the two forms of vorticity for a location in the mid-latitudes of the Northern Hemisphere.

Fig. 2.32. Illustration of the dependence of vorticity conservation on depth. To conserve both mass and angular momentum a cylinder of fluid originally rotating at a rate w, and with radius r and depth h, must double its depth and halve its radius if its rotation rate doubles.
Rangkaian Listrik Campuran

D is the depth of the body of water and Z is the relative vorticity. The latter, given mathematically by dv d u dx dy

is positive for anticlockwise rotation and negative for clockwise.

As the water in a Northern Hemisphere sub-tropical gyre rotates it must gain positive relative vorticity while moving south (as f decreases) and negative relative vorticity going north. In addition the water is subject to frictional retardation at the east and west coasts. The horizontal shear that this imparts to the flow results in an input of positive relative vorticity in the west and negative in the east (see Fig. 2.33). Thus, if the gyre was symmetrical, the negative vorticity added by the wind might be balanced by the positive vorticity contributed at the east coast, but on the western side of the basin there would be an excess of negative vorticity leading to an acceleration of the rotation. For balance the positive vorticity given by friction in the west needs to be high. This is achieved by a strong western boundary current trapped in a narrow zone near the coast. The Gulf Stream, in the North Atlantic, is such a current.

These strong boundary currents are important components of the climate system as they transfer large quantities of warm water polewards very rapidly. The Gulf Stream, for example, has a mean flow of about 100 x 106 m3s-1. If the Gulf Stream off the northern United States were only 2°C warmer than the global zonally averaged sea temperature for 40°N then this would represent a poleward energy flow of 1015 Js-1, approximately 20% of the total northward

Fig. 2.33. Pictorial representation of the various contributions to the vorticity of (a) a symmetric sub-tropical gyre and (b) a strongly asymmetric gyre with an intensified western boundary current. The flow in (b), in contrast to that in (a), enables a vorticity balance to be attained, so that the gyre's rotation does not accelerate indefinitely. [Fig. 4.13 of Open University Course Team (1989). Reproduced with permission of

Butterworth-Heinemann Ltd.]

Vorticity Open University 1989

heat flow in the atmosphere and ocean combined at these latitudes (this is actually an under-estimate of the local temperature anomaly in the Gulf Stream). The western boundary currents are as important in the ocean for poleward heat transport as mid-latitude depressions are in the atmosphere, although the latter probably contribute more than half to the global total of this transport. Any change in gyre circulation therefore has a significant impact on the climate system.

The boundary currents exhibit another feature, particularly after separating from the coast. The strong shear in flow across the current, and the thermal contrast between the warm sub-tropical waters equatorwards and the cold waters polewards, leads to instability. The current meanders, occasionally leaving intrusions of warm or cold water isolated in the cold or warm regions respectively. This process assists the diffusion of heat polewards, but also has implications for biological productivity, as we will discover in §2.11.6.

It takes a month or two for water to be transported in this boundary current from the tropics to the mid-latitudes, but numerical models suggest full circulation requires 15 years or more. Ekman transport suggests that during this circulation there is a net movement of water towards the centre of the gyre, as noted in §2.10.1. The accompanying doming of the surface, while only a few centimetres, maintains, and is maintained by, the geostrophic gyre circulation. The converging water forms a new water mass at intermediate depths in the ocean.

Poleward of the sub-tropical gyres the atmospheric circulation imparts positive relative vorticity to the ocean, as the surface winds swing from westerly to easterly. A sub-polar, cyclonic gyre would therefore be expected. Vorticity arguments imply that there should be strong western boundary currents in such a gyre as well. No such gyre exists in any of the Southern Hemisphere oceans, because of the circumpolar ocean link at 60°S (see §2.11.6). In the Northern Hemisphere there are weak gyres, heavily constricted by northern coasts, in both the Pacific and Atlantic Oceans. Their western boundary currents are consequently rather weak.

2.11.2 Coastal upwelling

The sub-tropical anticyclones produce equatorward winds along the eastern shores of the oceans, as shown in Fig. 1.6. The Ekman transport associated with these winds is off-shore. To balance this mass transport, cold water episodically upwells from below the surface, bringing high concentrations of nitrates and phosphates to the euphotic zone. These regions of coastal upwelling are therefore very biologically productive, as such chemicals are the basic food source for plankton (see Chapter 4). It is important to note that such regions are not generally found on western coasts of the ocean basins because of the asymmetry of the position of the anticyclones producing the wind forcing. The reasons for this will be pursued in Chapter 5.

The strongest upwelling occurs off the west coast of northern and southern Africa, and off the west coast of South America. The abundant plankton provide a base for the food chain in these waters, leading to rich fisheries. The upwelling zone off South America is particularly known for its anchovy fishing. This region

Fig. 2.34. Illustration of the surface transport due to the off-shore Ekman transport in regions of coastal upwelling in the Northern Hemisphere. [Fig. 4.30c of Open surlace transport current because resultant surface gradient i transport sea-surface,

University Course Team (1989). Reproduced with permission of surlace transport

Fig. 2.34. Illustration of the surface transport due to the off-shore Ekman transport in regions of coastal upwelling in the Northern Hemisphere. [Fig. 4.30c of Open

University Course Team (1989). Reproduced with permission of geoslioptitc current because resultant surface gradient i transport sea-surface, is, however, subject to occasional periods of dramatic reduction in the upwelling as a component of the climatic disturbances associated with El Nino: a coupling of the ocean and atmosphere causing the climate of the tropical Pacific basin, and further afield, to be highly anomalous for periods of a year or more. El Nino will be discussed in detail in Chapter 5.

Upwelled water, being colder and denser than that moved westwards by Ekman transport, does not quite replace the volume of the lost water. A slope is therefore set up off-shore, as shown in Fig. 2.34. This slope creates a pressure gradient towards the shore, with a consequent geostrophic current aligned alongshore with the wind direction. The observed surface ocean current is therefore a combination of the off-shore Ekman transport and this alongshore geostrophic flow; the result is at an angle to the coast, as in Fig. 2.34.

2.11.3 The tropical surface circulation

The wind systems of the two hemispheres meet in the tropics. We saw in §1.2 that the easterly winds on the equatorward side of the sub-tropical anticyclones converge in a semi-continuous line called the Inter-Tropical Convergence Zone, or ITCZ. The converging flow forces air to rise, creating the tendrils of massive thunderstorms seen in satellite photographs such as Fig. 2.35. This narrow band of storms moves in response to the seasons, tending to be in the summer hemisphere. Fig. 2.35 shows a clear dependence of the positioning of the ITCZ on the land/sea distribution, with heated land masses, such as central Africa, pulling the ITCZ polewards of its latitude over the neighbouring oceans.

The junction of the hemispheric Trade wind systems in the vicinity of the equator leads to interesting possibilities for the dynamical interaction of the ocean and atmosphere. This interaction is illustrated in Fig. 2.36. As the ITCZ tends to form in the summer hemisphere the Trade winds from the winter hemisphere must cross the equator. This circulation is subject to oppositely directed, diverging, Coriolis forces, on either side of the equator. The ocean circulation is also subject to these forces, and the resulting Ekman transport

Fig. 2.35. Infra-red image from the geostationary Meteostat satellite, taken at 1030 GMT on 11 April 1995. Note the vigorous convection over central Africa and the set of storms over the tropical Atlantic which are part of the ITCZ. [Supplied by the Department of Meteorology at the University of Edinburgh.]

Fig. 2.36. Schematic of the various currents in the tropical oceans and their connection to surface divergence or convergence and the wind field. [Fig. 5.1a from Open University Course Team (1989). Reproduced with permission of

Butterworth-Heinemann Ltd.]

Fig. 2.35. Infra-red image from the geostationary Meteostat satellite, taken at 1030 GMT on 11 April 1995. Note the vigorous convection over central Africa and the set of storms over the tropical Atlantic which are part of the ITCZ. [Supplied by the Department of Meteorology at the University of Edinburgh.]

causes divergence of surface waters from the equator. A narrow strip of ocean along the equator will therefore need upwelling to compensate for this Ekman flow. This zone is clearly seen as an equatorial tongue of cold water in the sea surface temperature pattern of Fig. 2.37, particularly in the Pacific Ocean. The horizontal surface circulation at the equator itself mirrors the westward

Fig. 2.37. Average sea surface temperature over the tropical oceans during January. Contour interval is 2°C.

wind stress, as the Coriolis force vanishes here. The steady winds can induce strong currents of more than 1 ms-1. The carbon exchange between ocean and atmosphere is also distinctly anomalous along the equator, as will be discussed in Chapter 3.

Polewards of the equator the main atmospheric, and oceanic, circulation consists of the equatorward side of the sub-tropical anticyclones, or gyres. The flow in both media is therefore predominantly westward. In the ocean this creates what are known as the North and South Equatorial Currents, shown in Fig. 1.15. This westward flow, coupled with the weakness of the Coriolis force very close to the equator, creates a gradient in the sea surface across the ocean basin. The resulting eastward pressure gradient opposes the westward flow, generated by the equatorial easterly winds, which acts to support the sea surface gradient. The upper ocean circulation pattern is balanced by water flowing below the surface eastwards, and towards the equator, at a depth of about 100 m. This converges to form the eastward flowing, Equatorial Undercurrent. This can also reach high speeds of well over 1 ms-1, the strongest flow being tightly constrained near the equator in most circumstances.

The equatorial oceans are thus regions of strong gradients and strong currents. Near the ITCZ itself yet another dynamic interaction occurs, as shown in Fig. 2.36 for a Northern Hemisphere summer. Here we find a region of light winds - the Doldrums - as the Trade wind systems mix. The Ekman transport produced by the summer hemisphere Trades, which have not crossed the equator, is away, and hence divergent, from the Doldrums. The winter hemisphere Trades, having crossed the equator, also produce an Ekman transport away from the equator. This transport, however, causes a convergence of water on the equatorward side of the Doldrums. The combination of this divergence and convergence creates a sea surface slope across the Doldrums. The resulting pressure gradient is polewards, producing a geostrophic flow to the east: the Equatorial Counter-Current.

The upper ocean dynamics illustrated schematically in Fig. 2.36 are theoretical in nature. In the central and eastern Pacific the ITCZ remains north of the equator throughout the year, while in the west the ITCZ breaks dramatically to the south in the southern summer. This complicates the current pattern considerably in the west Pacific. The Atlantic sector, by contrast, is closer to the theoretical picture, with the ITCZ moving substantially between the seasons.

The strength of the Southern Hemisphere Trades here can vary substantially with latitude at certain seasons, giving rise to a weak counter-current south of the equator.

The Indian Ocean sector has different behaviour again. The southern summer has a pattern strongly reminiscent of the theoretical description. However, in the northern summer the strong heating of the Eurasian continent dramatically accentuates the northward migration of the ITCZ. This leads to a radically different ocean-atmosphere interaction which we discuss in the next section.

2.11.4 The Indian Ocean monsoonal circulation

Monsoonal climates occur wherever there is a strong seasonal contrast between land and ocean temperatures. Their driving mechanism is the same as for a sea breeze, but on a gigantic scale. We will therefore start by examining why sea breezes occur.

During daylight hours the land surface is heated by the Sun. The warmed earth then heats the air above it, making it less dense than the surrounding atmosphere and causing it to rise. Near a coast where the sea surface temperature is cooler than the land, as is typical of summer, a pressure gradient will be established between the cooler, denser, air over the sea and the warmer, less dense, air over the land. A near-surface layer of air, up to a few hundred metres in height, will therefore flow from the sea to the land. Aloft, the rising air over the land, and the removal of air from near the surface over the sea, creates a pressure gradient in the opposite direction. A circulation cell is thus created.

Sea breezes usually only affect a narrow coastal strip a few kilometres in width, but can extend inland for several scores of kilometres in particularly favourable conditions. Their small geographical extent means that the Coriolis force has only a small influence on the circulation. During summer evenings the circulation cell can reverse, as the land cools relative to the sea, giving a land breeze. Both sea and land breezes can only occur when the prevailing large-scale, or synoptic, circulation is weak.

Monsoons are driven by similar forces. The Indian sub-continent and Tibetan plateau are strongly heated during the northern summer, becoming substantially warmer than the air temperatures at equivalent altitude over the Indian Ocean.6 This sets up a pressure gradient from the ocean towards India, but over much of the northern Indian Ocean. The net effect of this is shown in Fig. 1.6. The Southern Hemisphere Trade winds extend far north across the equator, the combination of the pressure gradient and Coriolis forces making the air flow in a great arc across the Arabian Sea. This flow feeds into the giant sea breeze system, fuelling the rising air over southern Asia and providing the moisture for the torrential rainfall that this produces. The strongest flow is in a narrow jet at a height of 1-2 km which curves over East Africa. This has the properties of an atmospheric western boundary current, with the East African Highlands

6 This is particularly true of the Tibetan plateau because of the much weaker attenuation of solar radiation at altitudes 4-5 km above sea level.

Fig. 2.38. Surface ocean circulation of the Indian Ocean in the northern summer. [Reprinted from G. L. Pickard and W. J. Emery, Descriptive Physical Oceanography, Copyright (1982), page 209, with kind permission from Elsevier Science Ltd., The Boulevard, Langford Lane, Kidlington OX5 1GB, UK.]

Fig. 2.38. Surface ocean circulation of the Indian Ocean in the northern summer. [Reprinted from G. L. Pickard and W. J. Emery, Descriptive Physical Oceanography, Copyright (1982), page 209, with kind permission from Elsevier Science Ltd., The Boulevard, Langford Lane, Kidlington OX5 1GB, UK.]

Wyrtki Jet

acting as the boundary. The land-sea contrast forces the summer northern movement of the ITCZ to be far greater than it would otherwise be. In the northern winter the equivalent 'land breeze' merely adds to the expected tropical circulation.

The Monsoon conditions force a number of special circulation features within the Indian Ocean itself. As the ITCZ crosses the equator during the equinoxal seasons the prevailing easterly winds weaken and reverse, giving westerlies. This produces a strong eastward-flowing jet at the surface of the ocean during these months, known as the Wyrtki jet after Klaus Wyrtki who led an international oceanographic expedition to the Indian Ocean in 1971 to investigate its circulation. This semi-annual variability in the equatorial flow means that the equatorial under-current is weak and not always evident in this ocean.

In the Arabian Sea the sweep of the summer Monsoon winds paralleling the African and Arabian coasts leads to upwelling, particularly off Somalia and Oman. The winds induce an off-shore Ekman transport, as discussed in §2.11.2. A series of small circulation cells, or eddies, are also established along the African coast north of the equator, within the strong coastal Somali current. The surface circulation of the Indian Ocean in summer is shown in Fig. 2.38.

The Indian Ocean is not the only location where monsoon conditions occur. Northern Australia has a strong monsoon circulation which is part of the climatic cycle of El Nino and will be discussed in Chapter 5. Southeast Asia and China also experience monsoonal flows.

A summer Monsoon occurs in west Africa, north of the equator, as well. The bulge of Africa that includes the Sahara, and extends out into the Atlantic ocean north of 10°N, is strongly heated during the summer. This draws the ITCZ unusually north in this region, bringing rains to the Sahel, which would otherwise be very arid. Disturbances to this pattern over the last 30 years have led to persistent drought affecting the Sahel. In §6.4.5 we will see that sea surface temperatures in the South Atlantic have decreased over this period, providing less energy and moisture to the southeasterly winds of the oceanic contribution to the ITCZ convergence.

2.11.5 The polar regions

One major oceanic impact on the atmosphere in the polar regions - reflection of solar radiation by pack ice - has already been discussed in various sections of Chapter 1. Where there is no pack ice, or frequent leads occur, the water is often warmer than the overlying air, providing a rich source of moisture, and hence latent heat, for the atmosphere. In sub-polar waters this frequently leads to the formation of active weather systems. The oceans are therefore major source regions for mid-latitude weather. This will be discussed in detail in §5.1.4.

The oceanic source of water vapour, coupled with the cold temperatures, leads to the polar oceans being particularly cloudy. Fog formation is also favoured. For example, sea smoke occurs where cold air is in contact with a warm sea that provides sufficient moisture to saturate the air. By contrast, in regions such as the Grand Banks off Newfoundland cold water flowing south from the Labrador Sea cools the air moving eastwards off a warming Canada in the spring to saturation. The abundance of latent heating potential in the atmosphere near the land-sea or sea-ice-sea boundary in polar regions makes these regions particularly subject to the development of intense storms, of smaller scale than regular depressions, called polar lows.

The atmosphere, in turn, has a strong impact on the polar oceans. In §1.3.2 we saw how strong winter cooling and storm-force winds cause instability in the water column in the seas north of Iceland. The surface waters mix with much of the water column, creating deep water which spreads out to fill much of the bottom layers of the world's oceans. This is shown schematically in Fig. 1.14. Such deep water formation is sporadic and localized. Very limited observational evidence suggests that it occurs in small regions, or chimneys, only a few tens of kilometres in diameter. Similar processes occur in the Gulf of Lyon, in the northwestern Mediterranean, contributing to the characterization of Mediterranean Water that eventually overflows into the North Atlantic.

The atmosphere also indirectly drives the formation of deep water originating from the Southern Ocean. This forms when sea water freezes to form sea-ice. The local release of salt, and consequent increase in surface water density, leads to relatively fresh but cold water sinking to the ocean floor to form Antarctic Bottom Water spreading as far north as 40°N in the Atlantic. The formation of this arm of the thermohaline circulation will be explored in more detail in §2.12. Chapter 5 will investigate the surprisingly important links between the climate system and the origin and circulation of bottom water.

Formation of intermediate depth water is mostly driven by the atmosphere, as discussed in §2.11.1. Some regions of the polar oceans also contribute to intermediate waters by large-scale cooling, as in the case of water in the Labrador Sea, or cooling and Ekman convergence in the Southern Ocean (Antarctic Intermediate Water).

The atmosphere also drives the surface circulation. The most pronounced polar circulation is the global Antarctic Circumpolar Current in the Southern Ocean, shown in Fig. 1.15. The driving force for this current is a continual, and strong, westerly wind stress. The Ekman transport is northwards, creating a slope downwards to the Antarctic coast in the isopycnals. The resulting offshore pressure gradient is balanced by the Coriolis force to produce the strong, geostrophically controlled, Circumpolar Current. At the surface itself, the wind induces a northward component to the flow, which leads to regions of strong latitudinal contrasts, or fronts.

The Circumpolar Current is strongly affected by the topography of the sea floor, and the intrusion into its zone of the land mass of South America. These boundary interactions produce regions of instability in the current, leading to copious formation of eddies which assist the mixing of water across the fronts.

2.11.6 Oceanic eddies

In several sub-sections of §2.11 we have encountered small circulation cells within the ocean called eddies. The meandering western boundary currents shed them during their unstable wandering. Seasonal eddies appear in the Somali current off eastern Africa. Eddies are created through instability in the Circumpolar Current. Throughout the world's oceans eddies can occur.

In the atmosphere the equivalent of eddies - the cyclones and anticyclones that make mid-latitude weather so variable and fascinating - are responsible for a major part of the poleward heat transport. In the ocean the eddies, while contributing to this climatic control, are very much the junior partners of the western boundary currents. The main reason for this ocean-atmosphere contrast is the relative size of the respective eddies. In the ocean eddies are scores of kilometres in diameter; in the atmosphere they can be more than a thousand. Atmospheric eddies also travel much faster, crossing the Atlantic in a few days, for example, while a Gulf Stream ring may wander only a few hundred kilometres over a season.

Where eddies may be most important in redistributing heat through the oceans is off South Africa. The western boundary current flowing south along the East African coast - the Agulhas Current - overshoots south and west of South Africa before being swept into the influence of the Circumpolar Current. Much of its water then turns east, as shown in Fig. 1.15. However, some of the warm water transported from the Indian Ocean is spun off into eddies that enter the Atlantic. A numerical model of the Southern Ocean - the Fine Resolution Antarctic Model - developed by scientists from the United Kingdom, suggests that this occurs quite regularly every 4-5 months, with the Indian Ocean water being recognizable far into the Atlantic. A snapshot of the circulation and temperature field from this model at 120 m is shown in Fig. 2.39.

Arnold Gordon, of the Lamont-Doherty Geological Laboratory near New York, has proposed a Conveyor Belt model for global oceanic heat transport and water circulation, where water from the Pacific leaks through the Indonesian Archipelago, crosses the Indian Ocean and, via these eddies from the Agulhas Current, flows into the Atlantic. This transport is hypothesized to close the ocean circulation, the other half of the circuit being the deep water that flows slowly from the Atlantic to the Pacific. Gordon's model is illustrated in Fig. 1.14. In addition to providing a surface flux of water to the Atlantic, these eddies from the Agulhas Current also carry heat. The net heat transport of the Atlantic Ocean is estimated to be northwards at all latitudes, even south of the equator. The apparent anomaly in the South Atlantic (where heat would be expected to be transported polewards) could be explained by this Conveyor Belt theory.

Longitude

Fig. 2.39. A snapshot of the velocity field at a depth of 120 m off South Africa from the Fine Resolution Antarctic Model. Note the strong Agulhas Current flowing south along the east coast of Africa and the eddies of approximately 200 km diameter spun off into the Atlantic. [Picture courtesy of David Stevens.]

Longitude

Consistent with this theory is the observation that heat flow in the Indian Ocean is everywhere southwards.

Eddies can also be biologically distinctive, compared with surrounding waters. Warm core eddies tend to be biologically poor, either with extensive down-welling or deriving from nutrient-poor waters. In either case the result is a limited nutrient supply. Cold core eddies, by contrast, are very productive as either there is upwelling, and so a continuous supply of nutrients from below the euphotic zone, or such eddies bring water from a nutrient-rich region. In Chapters 3 and 4 we shall see that this may have implications for air-sea fluxes of various gases.

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