Above the thermocline a Vertical

The major atmosphere-ocean interactive processes (Figure 7.26) involve heat exchanges, evaporation, density changes and wind shear. The effect of these processes is to produce a vertical oceanic layering that is of great climatic significance:

1 At the ocean surface, winds produce a thermally mixed surface layer averaging a few tens of metres deep poleward of latitude 60°, 400 m at latitude 40° and 100 to 200 m at the equator.

2 Below the relatively warm mixed layer is the thermocline, a layer in which temperature decreases and density increases (the pycnocline) markedly with depth. The thermocline layer, within which stable stratification tends to inhibit vertical mixing, acts as a barrier between the warmer surface water and the colder deep-layer water. In the open ocean between latitudes 60° north and south the thermocline layer extends from depths of about 200 m to a maximum

Meteorological And Oceanographic Process
Figure 7.26 Generalized depiction of the major atmosphere-ocean interaction processes. The sea ice thickness is not to scale. Source: Modified from NASA (n.d.). Courtesy of NASA.

of 1000 m (at the equator from about 200 to 800 m; at 40° latitude from about 400 to about 1100 m). Poleward of 60° latitude, the colder deep-layer water approaches the surface. The location of the steepest temperature gradient is termed the permanent thermocline, which has a dynamically inhibiting effect in the ocean similar to that of a major inversion in the atmosphere. However, heat exchanges take place between the oceans and the atmosphere by turbulent mixing above the permanent thermocline, as well as by upwelling and downwelling.

During spring and summer in the mid-latitudes, accentuated surface heating leads to the development of a seasonal thermocline occurring at depths of 50 to 100 m. Surface cooling and wind mixing tend to destroy this layer in autumn and winter.

Below the thermocline layer is a deep layer of cold, dense water. Within this, water movements are mainly driven by density variations, commonly due to salinity differences (i.e. having a thermohaline mechanism).

In terms also of circulation the ocean may be viewed as consisting of a large number of layers: the topmost subject to wind stress, the next layer down to frictional drag by the layer above, and so on; all layers being acted on by the Coriolis force. The surface water tends to be deflected to the right (in the northern hemisphere) by an angle averaging some 45° from the surface wind direction and moving at about 3 per cent of its velocity. This deflection increases with depth as the friction-driven velocity of the current decreases exponentially (Figure 7.27). On the equator, where there is no Coriolis force, the surface water moves in the same direction as the surface wind. This theoretical Ekman spiral was developed under assumptions of idealized ocean depth, wind constancy, uniform water viscosity and constant water pressure at a given depth. This is seldom the case in reality, and under most oceanic conditions

Ekman Spiral Atmosphere

Figure 1.21 The Ekman ocean current pattern in the northern hemisphere. Compare Figure 6.5.

Source: Bearman (1989). Copyright © Butterworth-Heinemann, Oxford.

Many large-scale characteristics of ocean dynamics resemble features of the atmosphere. These include: the general circulation, major oceanic gyres (similar to atmospheric subtropical high-pressure cells), major jet-like streams such as sections of the Gulf Stream (see Figure 7.29), large-scale areas of subsidence and uplift, the stabilizing layer of the permanent thermocline, boundary layer effects, frontal discontinuities created by temperature and density contrasts, and water mass ('mode water') regions.

Mesoscale characteristics that have atmospheric analogues are oceanic cyclonic and anticyclonic eddies, current meanders, cast-off ring vortices, jet filaments, and circulations produced by irregularities in the north equatorial current.

Figure 1.21 The Ekman ocean current pattern in the northern hemisphere. Compare Figure 6.5.

Source: Bearman (1989). Copyright © Butterworth-Heinemann, Oxford.

the thickness of the wind-driven Ekman layer is about 100 to 200 m. North (south) of 30°N, the westerly (easterly) winds create a southward (northward) transport of water in the Ekman layer giving rise to a convergence and sinking of water around 30°N, referred to as Ekman pumping.

b Horizontal

(1) General

Comparisons can be made between the structure and dynamics of the oceans and the atmosphere in respect of their behaviour above the permanent thermocline and below the tropopause - their two most significant stabilizing boundaries. Within these two zones, fluidlike circulations are maintained by meridional thermal energy gradients, dominantly directed poleward (Figure 7. 28), and acted upon by the Coriolis force. Prior to the 1970s oceanography was studied in a coarsely averaged spatial-temporal framework similar to that applied in classical climatology. Now, however, its similarities with modern meteorology are apparent. The major differences in behaviour between the oceans and the atmosphere derive from the greater density and viscosity of ocean waters and the much greater frictional constraints placed on their global movement.

(2) Macroscale

The most obvious feature of the surface oceanic circulation is the control exercised over it by the low-level planetary wind circulation, especially by the subtropical oceanic high-pressure cells and the westerlies. The oceanic circulation also displays seasonal reversals of flow in the monsoonal regions of the northern Indian Ocean, off East Africa and off northern Australia (see Figure 7.29). As water moves meridionally, the conservation of angular momentum implies changes in relative vorticity (see pp. 119 and 140), with poleward-moving currents acquiring anticyclonic vorticity and equatorward-moving currents acquiring cyclonic vorticity.

The more or less symmetrical atmospheric subtropical high-pressure cells produce oceanic gyres with centres displaced towards the west sides of the oceans in the northern hemisphere. The gyres in the southern hemisphere are more symmetrically located than those in the northern, due possibly to their connection with the powerful west wind drift. This results, for example, in the Brazil current being not much stronger than the Benguela current. The most powerful southern hemisphere current, the Agulhas, possesses nothing like the jet-like character of its northern counterparts.

Equatorward of the subtropical high-pressure cells, the persistent trade winds generate the broad north and south equatorial currents (see Figure 7.29). On the western sides of the oceans, most of this water swings poleward with the airflow and thereafter comes increasingly under the influence of the Coriolis deflection and of the anticyclonic vorticity effect. However, some

Figure 7.28 Mean annual meridional heat transport (I0l5W) in the Pacific, Atlantic and Indian Oceans, respectively (delineated by the dashed lines). The latitudes of maximum transport are indicated.

Source: Hastenrath (1980), by permission of the American Meteorological Society.

water tends to pile up near the equator on the western sides of oceans, partly because here the Ekman effect is virtually absent, with little poleward deflection and no reverse current at depth. To this is added some of the water that is displaced northward into the equatorial zone by the especially active subtropical high-pressure circulations of the southern hemisphere. This accumulated water flows back eastward down the hydraulic gradient as compensating narrow-surface equatorial counter-currents, unimpeded by the weak surface winds. Near the equator in the Pacific Ocean, upwelling raises the thermocline to only 50 to 100 m depth, and within this layer there exist thin, jet-like equatorial undercurrents flowing eastwards (under hydraulic gradients) at a speed of 1 to 1.5 m s-1.

As the circulations swing poleward around the western margins of the oceanic subtropical high-pressure cells, there is the tendency for water to pile up against the continents, giving, for example, an appreciably higher sea-level in the Gulf of Mexico than along the Atlantic coast of the United States. The accu mulated water cannot escape by sinking because of its relatively high temperature and resulting vertical stability. Consequently, it continues poleward driven by the dominant surface airflow, augmented by the geostrophic force acting at right-angles to the ocean surface slope. Through this movement, the current gains anticyclonic vorticity, reinforcing the similar tendency imparted by the winds, leading to relatively narrow currents of high velocity (for example, the Kuroshio, Brazil, Mozambique-Agulhas and, to a lesser degree, the East Australian current). In the North Atlantic, the configuration of the Caribbean Sea and Gulf of Mexico especially favours this pile-up of water, which is released poleward through the Florida Straits as the narrow and fast Gulf Stream (Figure 7.30). These poleward currents are opposed both by their friction with the nearby continental margins and by energy losses due to turbulent diffusion, such as those accompanying the formation and cutting off of meanders in the Gulf Stream. These poleward western boundary currents (e.g. the Gulf Stream and the Kuroshio current) are

Figure 7.29 The general ocean current circulation in January. This holds broadly for the year, except that in the northern summer some of the circulation in the northern Indian Ocean is reversed by the monsoonal airflow. The shaded areas show mean annual anomalies of ocean surface temperatures (°C) of greater than +5°C and less than -3°C.

Figure 7.29 The general ocean current circulation in January. This holds broadly for the year, except that in the northern summer some of the circulation in the northern Indian Ocean is reversed by the monsoonal airflow. The shaded areas show mean annual anomalies of ocean surface temperatures (°C) of greater than +5°C and less than -3°C.

Sources: US Naval Oceanographic Office and Niiler (1992). Courtesy of US Naval Oceanographic Office.

approximately 100 km wide and reach surface velocities greater than 2 m s-1. This contrasts with the slower, wider and more diffuse eastern boundary currents such as the Canary and California (approximately 1000 km wide with surface velocities generally less than 0.25 m s-1). The northward-flowing Gulf Stream causes a heat flux of 1.2 X1015 W, 75 per cent of which is lost to the atmosphere and 25 per cent in heating the Greenland-Norwegian seas area. On the poleward sides of the subtropical high-pressure cells westerly currents dominate, and where they are unimpeded by landmasses in the southern hemisphere they form the broad and swift west wind drift. This strong current, driven by unimpeded winds, occurs within the zone 50 to 65°S and is associated with a southward-sloping ocean surface generating a geostrophic force, which intensifies the flow. Within the west wind drift, the action of the Coriolis force produces a convergence zone at about 50°S marked by westerly submarine jet streams reaching velocities of 0.5 to 1 m s-1. South ofthe west wind drift, the Antarctic divergence with rising water is formed between it and the east wind drift closer to Antarctica. In the northern hemisphere, a great deal of the eastward-moving current in the Atlantic swings northward, leading to anomalously very high sea temperatures, and is compensated for by a southward flow of cold Arctic water at depth. However, more than half of the water mass comprising the North Atlantic current, and almost all that of the North Pacific current, swings south around the east sides of the subtropical high-pressure cells, forming the Canary and California currents. Their southern-hemisphere equivalents are the

Figure 7.30 Schematic map of the western North Atlantic showing the major types of ocean surface circulation. Source: From Tolmazin (1994) Copyright © Chapman and Hall.

Benguela, Humboldt (or Peru) and West Australian currents (Figure 7.29).

Ocean fronts are associated particularly with the poleward-margins of the western boundary currents. Temperature gradients can be 10°C over 50 km horizontally at the surface and weak gradients are distinguishable to several thousand metres' depth. Fronts also form between shelf water and deeper waters where there is convergence and downwelling.

Another large-scale feature of ocean circulation, analogous to the atmosphere, is the Rossby wave. These large oscillations have horizontal wavelengths of 100s-1000s km and periods of tens of days. They develop in the open ocean of mid-latitudes in eastward-flowing currents. In equatorial, westward-flowing currents, there are faster, very long wavelength Kelvin waves (analogous to those in the lower stratosphere)

(3) Mesoscale

Mesoscale eddies and rings in the upper ocean are generated by a number of mechanisms, sometimes by atmospheric convergence or divergence, or by the casting off of vortices by currents such as the Gulf Stream where it becomes unsteady at around 65°W (Figure 7.30). Oceanographic eddies occur on the scale of 50 to 400 km in diameter and are analogous to atmospheric low- and high-pressure systems. Ocean mesoscale systems are much smaller than atmospheric depressions (which average about 1000 km in diameter), travel much slower (a few kilometres per day, compared with about 1000 km per day for a depression) and persist from one to several months (compared with a depression life of about a week). Their maximum rotational velocities occur at a depth of about 150 m, but the vortex circulation is observed throughout the thermocline (ca. 1000 m depth). Some eddies move parallel to the main

Drainage Hilly Areas

Figure 7.31 Schematic illustration of mechanisms that cause ocean upwelling. The large arrows indicate the dominant wind direction and the small arrows the currents. (A) The effects of a persistent offshore wind. (B) Divergent surface currents. (C) Deep-current shoaling. (D) Ekman motion with coastal blocking (northern hemisphere case).

Source: Partly modified after Stowe (1983) Copyright 1983 © John Wiley & Sons, Inc. Reproduced by permission.

Figure 7.31 Schematic illustration of mechanisms that cause ocean upwelling. The large arrows indicate the dominant wind direction and the small arrows the currents. (A) The effects of a persistent offshore wind. (B) Divergent surface currents. (C) Deep-current shoaling. (D) Ekman motion with coastal blocking (northern hemisphere case).

Source: Partly modified after Stowe (1983) Copyright 1983 © John Wiley & Sons, Inc. Reproduced by permission.

flow direction, but many move irregularly equatorward or poleward. In the North Atlantic, this produces a 'synoptic-like' situation in which up to 50 per cent of the area may be occupied by mesoscale eddies (see Plate B). Cold-core cyclonic rings (100 to 300-km diameter) are about twice as numerous as warm-core anticyclonic eddies (100-km diameter), and have a maximum rotational velocity of about 1.5 m s-1. About ten cold-core rings are formed annually by the Gulf Stream and may occupy 10 per cent of the Sargasso Sea.

2 Deep ocean water interactions a Upwelling

In contrast with the currents on the west sides of the oceans, equatorward-flowing eastern currents acquire cyclonic vorticity, which is in opposition to the anticyclonic wind tendency, leading to relatively broad flows of low velocity. In addition, the deflection due to the Ekman effect causes the surface water to move westward away from the coasts, leading to replacement by the upwelling of cold water from depths of 100 to 300 m (Figure 7.31 A, D). Average rates of upwelling are low (1 to 2 m/day), being about the same as the offshore surface current velocities with which they are balanced. The rate of upwelling therefore varies with the surface wind stress. As the latter is proportional to the square of the wind speed, small changes in wind velocity can lead to marked variations in rates of upwelling. Although the band of upwelling is of limited width (about 200 km for the Benguela current), the Ekman effect spreads this cold water westward. On the poleward margins of these cold-water coasts, the meridional swing of the wind belts imparts a strong seasonality to the upwelling; the California current upwelling, for example (Plate 16), is particularly well marked during the period March to July.

A major region of deep-water upwelling is along the West Coast of South America (Figure 11.52) where there is a narrow 20-km-wide shelf and offshore easterly winds. Transport is offshore in the upper 20 m but onshore at 30 to 80 m depth. This pattern is forced by the offshore airflow normally associated with the large-scale convective Walker cell (see Chapters 7C.1 and 11G) linking Southeast Asia-Indonesia with the eastern South Pacific. Every two to ten years or so this pressure difference is reversed, producing an El Niño event with weakening trade winds and a pulse of warm surface water spreading eastward over the South Pacific, raising local sea surface temperatures by several degrees.

Coastal upwelling is also caused by less important mechanisms such as surface current divergence or the effect of the ocean bottom configuration (see Figure 7.31 B, C).

b Deep ocean circulation

Above the permanent thermocline the ocean circulation is mainly wind driven, while in the deep ocean it is driven by density gradients due to salinity and temperature differences - a thermohaline circulation. These

Ocean Thermohaline Circulation

Figure 7.32 The deep ocean thermohaline circulation system leading to Broecker's concept of the oceanic conveyor belt.

Source: Kerr (1988). Reprinted with permission from Science 239, Fig. 259. Copyright © 1988 American Association for the Advancement of Science.

Figure 7.32 The deep ocean thermohaline circulation system leading to Broecker's concept of the oceanic conveyor belt.

Source: Kerr (1988). Reprinted with permission from Science 239, Fig. 259. Copyright © 1988 American Association for the Advancement of Science.

differences are mostly produced by surface processes, which feed cold, saline water to the deep ocean basins in compensation for the deep water delivered to the surface by upwelling. Although upwelling occurs chiefly in narrow coastal locations, subsidence takes place largely in two broad ocean regions - the northern North Atlantic and around parts of Antarctica (e.g. the Weddell Sea).

In the North Atlantic, particularly in winter, heating and evaporation produce warm, saline water which flows northward both in the near-surface Gulf Stream-North Atlantic current and at intermediate depths of around 800 m. In the Norwegian and Greenland seas, its density is enhanced by further evaporation due to high winds, by the formation of sea ice, which expels brine during ice growth, and by cooling. Exposed to evaporation and to the chill high-latitude airmasses, the surface water cools from about 10° to 2°C, releasing immense amounts of heat into the atmosphere, supplementing solar insolation there by some 25 to 30 per cent and heating western Europe.

The resulting dense high-latitude water, equivalent in volume to about twenty times the combined discharge of all the world's rivers, sinks to the bottom of the North Atlantic and fuels a southward-flowing density current, which forms part of a global deep-water conveyor belt (Figure 7.32). This broad, slow and diffuse flow, occurring at depths of greater than 1500 m, is augmented in the South Atlantic/circum-Antarctic/Weddell Sea region by more cold, saline, dense subsiding water. The conveyor belt then flows eastward under the Coriolis influence, turning north into the Indian and, especially, the Pacific Ocean. The time taken for the conveyor belt circulation to move from the North Atlantic to the North Pacific has been estimated at 500 to 1000 years. In the Pacific and Indian Oceans, a decrease of salinity due to water mixing causes the conveyor belt to rise and to form a less deep return flow to the Atlantic, the whole global circulation occupying some 1500 years or so. An important aspect of this conveyor belt flow is that the western Pacific Ocean contains a deep source of warm summer water (29°C) (Figure 7.33). This heat

165°W 120°W 60° W 0° 60°E 120°E 180°E

Figure 7.33 Mean ocean-surface temperatures (°C) for January and July. Comparison of these maps with those of mean sea-level air temperatures (Figure 3.11) shows similarities during the summer but a significant difference in the winter.

Figure 7.33 Mean ocean-surface temperatures (°C) for January and July. Comparison of these maps with those of mean sea-level air temperatures (Figure 3.11) shows similarities during the summer but a significant difference in the winter.

Source: Reprinted from Bottomley et al. (1990), by permission of the Meteorological Office. Crown copyright ©.

differential with the eastern Pacific assists the highphase Walker circulation (see Figure 7.22A).

The thermal significance of the conveyor belt implies that any change in it may promote climatic changes operating on time scales of several hundred or thousand years. However, it has been argued that any impediment to the rise of deep conveyor belt water might cause ocean surface temperatures to drop by 6°C within thirty years at latitudes north of 60°N. Changes to the conveyor belt circulation could be initiated by lowering the salinity of the surface water of the North Atlantic; for example, through increased precipitation, ice melting, or fresh-water inflow. The complex mechanisms involved in the deep ocean conveyor belt are still poorly understood.

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