Dynamic topography
5.2 MONSOONAL CIRCULATION
5.2.2 THE CURRENT SYSTEM OF THE INDIAN OCEAN
As you might expect, the surface circulation of the northern Indian Ocean changes seasonally in response to the monsoons, but most resembles that of the other two oceans in the northern winter, during the North-East Monsoon (cf. Section 5.2.1). At this time of year, both a North and a South Equatorial Current are present, as well as an Equatorial Counter-Current (Figure 5.12(a),(b)). In the northern summer, by contrast, the flow in the North Equatorial Current reverses and combines with a weakened Equatorial Counter-Current to form the South-West Monsoon Current (Figure 5.12(d),(e)). The South Equatorial Current is still present, although its flow is not as strong as during the North-East Monsoon.
Figure 5.12 Surface currents in the northern Indian Ocean, as deduced from ships' drift data.
The North-East Monsoon is most fully established from January to March ((a),(b)), and the South-West Monsoon is most fully established during July-September ((d),(e)). The thicknesses of the lines indicate the relative intensities of the flows. For example, current speeds in the Equatorial Jet ((c),(f)) may reach 1.0-1.3 m s -1, but are mostly 0.3-0.7 m s -1 . The South Equatorial Current (SEC)is spread over the area between the two flow lines. The westward flow along the Equator within the South-West Monsoon Current, shown in (d), is discussed in Section 5.3.1. EAC is the East Arabian Current.
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Only for part of the year, not as a permanent feature. The existence of an Undercurrent depends on there being wind stress towards the west, to drive a westward surface current and hence cause a sea-surface slope up to the west; this would provide an eastward horizontal pressure gradient force to drive an eastward current below the wind-driven layer. In the equatorial Indian Ocean, the direction of the wind varies seasonally, but best fulfils the necessary conditions for an Undercurrent during the first few months of the year. The first direct observations of the Undercurrent were made during the International Indian Ocean Expedition (1962-65) at the end of the North- East Monsoon, when it was seen to be an ocean-wide feature; but this may not have been typical. In general, in the Indian Ocean the Undercurrent seems to be a stronger and more persistent flow in the western part of the ocean than in the central or eastern parts.
The changing pattern of surface currents in Figure 5.12 was deduced from ships' drift data collected by the UK Meteorological Office from log books of merchant vessels. Such information can provide reasonably accurate estimates of surface flow velocities in the Indian Ocean because, for most of the year, currents here are generally stronger than those in the Pacific and Atlantic Oceans.
It is an easm'ard Equatorial Jet, which is driven by westerly winds over the central equatorial ocean. Although the jet is detected in ships' drift data between April and June, and in October/November, it is possible that it is in fact a brief event, lasting perhaps only a month at a time.
When winds and currents along the Equator become westward, the sea- surface will (eventually) slope up to the west, and the thermocline slope down to the west; when winds and currents along the Equator are eastward, the sea-surface will (eventually) slope up to the east, and the thermocline slope down to the east.
The complexity of the surface circulation of the Indian Ocean, which contrasts with the relative simplicity of the gyral systems of the Pacific and Atlantic Oceans, is a result of the frequency and rapidity with which the overlying wind system changes. Wind speeds and directions change so fast that there is not always time for the upper ocean to adjust so that it is in equilibrium with the w i n d - as a result, during the relatively short inter- monsoon period, while the prevailing wind directions are changing dramatically, there is an increase in the large number of eddies, both cyclonic and anticyclonic and ranging in size from 100 to 1000 km across.
The ocean's response time - or, looked at another way, its ' m e m o r y ' - is many times longer than that of the atmosphere, and one of the most interesting questions that can be asked about the ocean circulation is:
How is it possible for the ocean to react as fast as it does?
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o ~. 9 9 ,.~ , rae ,.i,. t'~ ~'~ = ~ X >--~ ~ r _ .~=~< ~~ o~ 9 -,<6' -~ =- =The changes in the slopes of the sea-surface and thermocline along the Equator in the Indian Ocean, mentioned in Section 5.2.2, occur surprisingly fast. Exactly how fast the ocean can respond to seasonal changes in the wind has been studied in the simpler and steadier Atlantic, where the stress of the South-East Trade Winds across the equatorial ocean causes the sea- surface to slope up towards the west, and the thermocline to slope down (cf. Figure 5.3).
In the Atlantic, the South-East Trades are at their weakest during March- April and at their strongest during August-September. Figure 5.15 shows how the slopes of the sea-surface and thermocline across the Atlantic basin vary over the course of the year.
Figure 5.15 (a) Seasonal variations in the height of the sea-surface across the equatorial Atlantic.
(The curves appear jagged because they are based on mean values at particular longitudes.) (b) Seasonal variations in the depth of the 23 ~ isotherm across the equatorial Atlantic.
Figure 5.16 Examples of (a) a 'surface' long wave and (b) a long wave in the thermocline, here shown as a sharp boundary between less dense surface water (light greenish-blue) and more dense deeper water (darker blue). In (a), the surface ocean as a whole moves up and down, and isobaric and isopycnic surfaces remain parallel. Such waves are therefore described as 'barotropic'. In (b), the passage of the wave changes the vertical density
distribution, so that isopycnic surfaces are alternately compressed and separated. In addition, there are pressure variations over the surface of the density interface (the light greenish-blue/dark blue boundary) so that isobaric and isopycnic surfaces intersect; such waves are therefore described as 'baroclinic'.
163 Yes, they do, as far as the monthly mean values in Figure 5.15 allow us to tell. The important point about this is not the correlation itself but the fact that it indicates that the upper waters of the Atlantic Ocean respond ~,erv quickly indeed to changes in the overlying wind field.
This fast response time cannot be explained simply in terms of water being transported across the equatorial Atlantic and 'piling up" at the western boundary: the upper ocean as a whole must somehow adjust to the overlying wind and to the fact that there is a boundary along the western side of the ocean. In other words, the ocean in the central and eastern parts of the basin must in some sense 'know about' or "have felt' the western boundary. The way in which the mid-ocean receives information about the existence of a boundary is through perturbations or disturbances that travel through the ocean as pulses or waves. This is analogous to the way in which
~'e receive information concerning the world about us - through light waves or sound waves.
These wave-like disturbances not only enable water in mid-ocean to respond to the existence of coastal boundaries, they also transmit the effects of changes in the overlying wind field from one region of ocean to another, and do so much faster than would be possible simply through transportation of water in wind-driven currents. It is believed that the speed of the reversal of the Somali Current is one example of such an effect: as described in Section 5.2.2, the surface waters off Somalia flow fairly fast south-westwards during the North-East Monsoon, but can nevertheless be moving ~,erv fast in the opposite direction only a few months later. This dramatic 'switch" is thought to be possible because the upper ocean off eastern Africa 'feels' the effect of winds blowing in the central region, as well as being directly driven by local winds.
Various different types of waves may be generated in the ocean. We are all familiar with the wind waves which occur at the surface of the ocean, in which particles of water are displaced from their 'normal' or equilibrium position in a vertical direction and return under the influence of gravity.
The types of gravity waves that are of interest as far as ocean circulation is concerned have wavelengths which may be anything from tens to thousands of kilometres, periods ranging from days up to months, or even years, and vertical displacements which vary from centimetres up to tens of metres.
These vertical displacements may be more or less constant with depth (Figure 5.16(a)): alternatively, they may be greatest where there is a strong vertical density g r a d i e n t - in the thermocline, for example (Figure 5.16(b)).
In the first case, the vertical density distribution is not affected by the passage of the wave, and so if isobaric and isopycnic surfaces are initially parallel to one another, they will remain so (Figure 5.16(a)). For this reason, such waves are sometimes referred to as "barotropic waves': we will refer to them as 'surface waves" although it should be remembered that motion associated with them extends to significant depths, because their wavelengths are much greater than the depth of the ocean. In the second case, the vertical density distribution is affected by the passage of the waves, so that density surfaces are caused to intersect isobaric surfaces (Figure 5.16(b)): such waves are therefore referred to as 'baroclinic' waves.
Baroclinic waves generally have much larger amplitudes than surface (barotropic) waves.
Because these disturbances have very long wavelengths and periods, they are significantly affected by the Coriolis force. As a result, motion occurs in a h o r i z o n t a l as well as a vertical direction. This may be seen most clearly in the flow patterns associated with Kelvin waves and planetary or Rossby waves, the two classes of waves that are most important as far as ocean circulation is concerned.
A general feature of all wave motions, through water or any other medium, is that where the physical characteristics of the medium change with position, waves seeking to cross the line of change may be reflected, or in some other way deflected, so that they become trapped within a wave guide. Common examples of wave guides employed in communications technology are optical fibres and coaxial cables, both of which are used to carry information along a specified path. An example of a wave guide in the ocean is the sound channel, a depth zone where the velocity of sound in seawater is relatively low, and within which sound waves may be trapped by refraction (Section 4.3.3). Sounds emitted in the sound channel - by, for example, Sofar floats - are transmitted over distances of thousands of kilometres with relatively little attenuation.
Now imagine a parcel of water in the Northern Hemisphere, moving northwards with a coastal boundary on its right. The Coriolis force continually tends to deflect the parcel to the right, but because the coastal boundary is in the way, only limited deflection is possible. Water piles up against the boundary, giving rise to an offshore horizontal pressure gradient force; this keeps the parcel of water moving parallel to the coast, in a geostrophic current. Consequently, a coastal boundary constrains the way in which water can move in response to the forces acting on it. As a result, coasts may act as wave guides to the perturbations known as Kelvin waves (Figure 5.17(a)), which can travel as surface (barotropic) waves or as baroclinic waves.
In a Kelvin wave, perturbations of the sea-surface or of the thermocline propagate parallel to and close to the coast, as though unaffected by the Earth's rotation, because the Coriolis force directed towards the coast is opposed by a horizontal pressure gradient force that results from the slope of the sea-surface. Thus, a necessary condition for the propagation of Kelvin waves is that the horizontal pressure gradient force and Coriolis force act in opposition.
Kelvin waves are similar to surface wind waves in that the principal
maintaining force is gravity. Particle movement within the wave is such that the amplitude of the vertical displacement is greatest at the coast and decreases exponentially away from it, so that at any point and any time the Coriolis force balances the pressure gradient resulting from the slope of the sea-surface (or thermocline) (Figure 5.17(b)).
165
Figure 5.17 (a) Schematic diagram of a surface Kelvin wave in the Northern Hemisphere. In the case of a Kelvin wave in the thermocline, the thermocline would adopt the shape shown, while the sea-surface would take on a similar shape in 'mirror image', although to a much lesser extent.
(b) Diagram to show the horizontal balance of forces in a Kelvin wave, and the Rossby radius, L (which is discussed in the text).
(c) An equatorial 'double' Kelvin wave (here shown for a surface wave); the wave travels from west to east.
A Kelvin wave may be regarded as being 'trapped' within a certain distance of the coast, because by that distance its amplitude has significantly decayed away. This distance is known as the Rossby r a d i u s of d e f o r m a t i o n (L) and can be calculated from L =
c/f
where f is the Coriolis parameter and c is the wave speed. In mid-latitudes, the Rossby radius for a Kelvin wave with its maximum displacement in the thermocline is generally about 25 kin.Thus, in low latitudes, coastal Kelvin waves are not as closely trapped to the coast as they are in mid-latitudes.
The Rossby radius of deformation is not simply a measure of the degree to
which Kelvin waves are trapped. More generally, it is the distance that a wave with speed c can travel in time 1/[~ and it therefore provides a guide to the distance that a wave (i.e. a disturbance) can travel before being significantly affected by the Coriolis force. Because f is zero at the Equator and a maximum at the poles, Rossby radii decrease from infinity at the Equator to a minimum at the poles. The tendency for disturbances in current patterns to take on a curved or gyral character therefore increases with increasing distance from the Equator, and as a result, the higher the latitude, the smaller ('tighter') eddies, or other wave-like disturbances, tend to be. (This effect can be thought of as a manifestation of the increase in planetary vorticity with latitude, as illustrated in Figure 4.7(a).)
Almost everybody has experienced the effects of a Kelvin wave at first hand.
The twice-daily rise and fall of sea-level corresponding to high and low tide occurs in the form of coastal Kelvin waves, which progress anticlockwise round ocean basins (i.e. with the coast on the right) in the a m p h i d r o m i e systems of the Northern Hemisphere, and clockwise round basins in the Southern Hemisphere.
(Note: Tides and wind-generated waves are discussed elsewhere in this Series.) We have seen that coastal Kelvin waves may travel along coastal boundaries because the Coriolis force cannot play the same part in the balance of forces as it usually does. Along the Equator, the Coriolis force actually is zero and a similar effect arises, leading to the existence of an e q u a t o r i a l wave guide.
An equatorial Kelvin wave is like two parallel coastal Kelvin waves (one in each hemisphere) joining at the Equator, which they 'feel' as a boundary (Figure 5.17(c)). Like coastal Kelvin waves, equatorial Kelvin waves
propagate with the 'boundary' on the right in the Northern Hemisphere and on the left in the Southern Hemisphere. As a result, Kelvin waves can only propagate eastwards along the equatorial wave guide.
Surface equatorial Kelvin waves travel very fast, at about 200 m s -v. Their Rossby radius of deformation is about 2000 km, and so they can hardly be regarded as 'trapped' at all. However, this is not the case for Kelvin waves in the thermocline. They travel much more slowly, with c typically between 0.5 and 3.0 m s -l, and have Rossby radii of 100-250 km. They may be detectable at the surface, as sea-level is slightly raised above regions where the thermocline is depressed, and slightly depressed above regions where the thermocline is raised.
Kelvin waves in the thermocline can have dramatic effects, particularly in low latitudes where the mixed surface layer is thin. They may be generated by an abrupt change in the overlying wind field, as occurs for instance in the western Atlantic when the ITCZ moves northwards over the region and it comes under the influence of the South-East Trades. This causes a disturbance in the upper ocean (cf. Figure 5.18(a)), which travels eastwards along the equatorial wave guide as a double Kelvin wave in the thermocline (this takes about 4 - 6 weeks) and, on reaching the coast, splits into two coastal Kelvin waves, each travelling away from the Equator (Figure 5.18(b)). In the region of the disturbance where the thermocline bulges upward, colder, deeper water comes nearer to the surface.
By the time the wave reaches the coasts of Ghana and Ivory Coast (north of the Equator) and Gabon (south of the Equator), this cooler, sub-thermocline water is detectable at the surface, contributing to the 'seasonal coastal upwelling' off Abidjan and south of Cape Lopez, shown in light green in Figure 5.9.
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Now imagine a row of such water parcels in a current or airstream flowing along a line of latitude. If the flow is displaced polewards or equatorwards, horizontal oscillations of the kind described above may occur, so that the flow undulates about the original line of latitude (Figure 5.19(b)); these undulations are Rossby or planetary waves. In the ocean, the scale of such undulations is of the order of hundreds of kilometres; in the atmosphere, it varies from about 5000 to 20 000 km.
The large-scale undulations in the jet stream of the upper westerlies, shown in Figure 2.7, are atmospheric Rossby waves.
In both the atmosphere and the ocean, the overall effect of the clockwise and anticlockwise rotations associated with Rossby waves is to cause the wave-form- i.e. the undulations - to move westwards relative to the flow.
This is true even if the flow is itself moving eastwards, as is the case with the jet stream in the upper westerlies. However, flow velocities in the
Figure 5.19 (a) Diagram to show how in a Rossby wave the need to conserve potential vorticity (f + ()/D leads to a parcel of water oscillating about a line of latitude e while alternately gaining and losing relative vorticity (.
For details, see text.
(b) The path taken by a current or airstream affected by a Rossby wave. Note that the flow pattern is characterized by anticyclonic and cyclonic eddies, and that the wave-form moves westward relative to the current or airstream.
169 atmosphere may reach 100 m s -l and so Rossby waves in an airstream may move eastward relati~'e to the Earth, while still moving westward relative to the airstream. If the eastward motion of air in the airstream is approximately equal to the westward motion of the wave-form, stationary Rossby waves result. In the ocean, flow velocities rarely reach 1 m s -~ and so even in eastward-flowing currents, Rossby waves nearly always move westward relative to the Earth. Indeed, the Antarctic Circumpolar Current is the only current in which Rossby waves are carried eastward.
The way in which Kelvin and Rossby waves affect ocean circulation depends on the latitude. At middle and high latitudes, information about a change in the wind stress propagates mainly westwards, by means of Rossby waves, so the ocean near w'estern boundaries is affected by events in mid-ocean to a much greater extent than the ocean near eastern boundaries. By contrast, at low latitudes information can travel westwards by Rossby waves or eastwards by Kelvin waves in the equatorial wave guide. In addition, because of the equatorial wave guide, the upper ocean in low latitudes can respond to changing winds much faster than is possible away from the Equator. This is partly because the equatorial wave guide supports both Rossby and Kelvin waves, and partly because Rossby waves travel fastest there. For example, a Rossby wave can take as little as three months to travel west across the equatorial Pacific, whereas it could take ten years to cross the Pacific at 30 ~ N or 30 ~ S.
It would not be appropriate to go further into the details of either Rossby or Kelvin waves in this Volume. However, one of the most intriguing aspects of these waves is that when an equatorial Kelvin wave reaches the eastern boundary, it not only splits and travel polewards along the coast (as described in Section 5.3. l, for the tropical Atlantic), but may also be partially reflected as a Rossby wave. This can be seen in the computer- generated diagrams shown in Figure 5.20.
Figure 5.20 Computer-generated diagrams sllowang tile progress from mid-Pacific to the South American coast, of an internal equatorial Kelvin wave. The contour numbers may be regarded as either the depression of the thermocline in metres or the accompanying rise in sea-level in centimetres. The diagrams show the situation at successive monthly intervals. In (c), the equatorial Kelvin wave has split into two poleward-travelling coastal Kelvin waves.
Note that the coastal boundary has the effect of increasing the amplitude of the disturbance.
The equatorial Kelvin wave has also just been partially reflected as an equatorial Rossby wave, as can be seen by the circular contours which result from the rotatory motion associated with the wave. Because the two eddies are on either side of the Equator, both are anticyclonic and lead to topographic highs (H), although the northerly one is clockwise and the southerly one anticlockwise (cf. Figure 5.19(b)). (In (b) and (c), the small-scale waves in the contours are artefacts of the model.)
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171 El Nifio events are perturbations of the ocealt-atmosphere system. It is not known whether the perturbations originate in the atmosphere or the ocean, but for convenience we will start by considering what happens in the atmosphere during an El Nifio event. The prevailing winds over the equatorial Pacific are the South-East Trades. Their strength depends on the difference in surface atmospheric pressure between the subtropical high pressure region in the eastern South Pacific - where cool, dry air converges and s u b s i d e s - and the low pressure region over Indonesia- where warm, moist air rises, producing cumulonimbus clouds and heavy rainfall (Figure 5.21). During an E1 Nifio event, the Indonesian Low has anomalously high pressure (i.e. is a weaker low than usual) and moves eastwards into the central Pacific, while the South Pacific High becomes anomalously low. The South-East Trades weaken, and there are bursts of westerlies in the western Pacific.
The sea-surface slope will "collapse', so that both it and the thermocline become near-horizontal, enabling a considerable volume of warm mixed- layer water to move eastwards across the ocean. In the western Pacific, the collapse in the Trade Winds occurs abruptly, and so the resulting change in the upper o c e a n - a depression in the thermocline accompanied by a slight rise in sea-level (cf. Figure 5 . 1 8 ( a ) ) - propagates eastwards along the Equator as a pulse, or series of pulses, of Kelvin waves. At the eastern boundary, the equatorial Kelvin waves split into northward- and southward- travelling coastal Kelvin waves (cf. Figure 5.18(b)), as well as being partially reflected as Rossby waves (cf. Figure 5.20). The speed of these Kelvin waves has been calculated to be about 2.5 m s -~, but the bulge travels slightly faster than this because it is carried forward by flow in the Equatorial Undercurrent.
Figures 5.22 and 5.23 (overleaf) summarize the main differences between normal conditions in the Pacific basin, and conditions during an El Nifio event. As shown in part (a) of these Figures, the highest sea-surface temperatures of 28-29 ~ are normally found in the western ocean; during an El Nifio event, this area of exceptionally warm water moves into the central ocean, along with the vigorous convection of moist air usually associated with the Indonesian Low, and the Intertropical Convergence Zone shifts southwards and eastwards (Figure 5.23(b)).
A warm sea-surface leads to increased upward convection of moist air and, as discussed in Section 2.3.1, the increase in convection is particularly marked when sea-surface temperatures exceed -~28 ~ The eastward movement across the Pacific of the Indonesian Low, the shift in the position of the ITCZ, and the exceptionally warm surface water are therefore all intimately linked together. Thus, although the general eastward movement of warm mixed-layer water may be explained in terms of the slackening of the Trade Winds in response to a change in the strength and position of the Indonesian Low, this is clearly only part of the story.