The Sun The Wind And The

Energy2green Wind And Solar Power System

Wind Energy DIY Guide

Get Instant Access

The Sun and the Sea

When sunlight shines on the sea, what becomes of it? Some is reflected back into the sky, and the rest penetrates the surface to contribute its energy to the water; in a word, it is absorbed. This statement prompts two questions. What are the proportions reflected and absorbed? And what becomes of the absorbed energy?

Consider the first question. Recall the account in chapter 4 of the solar energy budget for the earth as a whole. It is believed (see fig. 4.1) that, averaging over all seasons, approximately 70 percent of all incoming sunlight is reflected back into space. This quantity—70 percent—is the earth's reflectivity or, to use the technical term, its albedo. Seventy percent is an average for the whole earth, but different surfaces have markedly different albedos. For example, the albedo of new-fallen snow is extremely high, up to 95 percent; that of calm water (this includes the ocean), is occasionally as low as 2 percent, rising to about 10

Incoming energy 100%

Outgoing energy

Incoming energy 100%

Figure 6.1. The fate of the short wave solar radiation absorbed by the sea (downward-pointing arrow on the left). The arrows pointing upward on the right show the proportions of incoming energy dissipated in various ways. See text for details.

Incoming sunlight (short wave)

Evaporation To space Conduction (long wave)

Figure 6.1. The fate of the short wave solar radiation absorbed by the sea (downward-pointing arrow on the left). The arrows pointing upward on the right show the proportions of incoming energy dissipated in various ways. See text for details.

percent for a rough sea; that of clouds is anywhere between 30 and 90 percent, depending on the type and thickness of the clouds.1

The low albedo of the ocean, ranging from 2 to 10 percent, doesn't mean, however, that the oceans always absorb from 90 to 98 percent of incoming solar radiation: far from it. The sunlight is often intercepted by strongly reflective clouds in the atmosphere before it ever reaches the ocean. Indeed, cloudiness probably has a stronger influence on an area's albedo than any other factor.2 Thus when you look over a calm, blue, sunlit sea, it is absorbing nearly all of the solar energy falling on it; this is especially true at low latitudes, where the sun's rays are nearly vertical at midday; but when you look at a gray, storm-tossed sea beneath a gray sky covered with thick clouds, the sea may be absorbing only 10 percent or less of the energy reaching the cloud tops.

Now for the fate of the absorbed solar energy, which is augmented to a negligible degree by heat from the earth's interior, and also by the frictional heat generated by the movements of seawater, especially surf. Figure 6.1 shows what happens to it.3 All the energy must ultimately be returned to space; if it were not, the oceans would heat up. Therefore the incoming energy balances the outgoing. More than half of it is disposed of by evaporative cooling; second in importance is heat loss in the form of long wave (infrared) radiation to space; last, a small amount is lost by conduction to the air in contact with the sea surface. Typical proportions are shown in the figure.

Some solar energy is absorbed only a short distance below the ocean surface, by living things—green plankton carrying on photosynthesis. This energy is subsequently converted to heat by living organisms as they breathe and by dead ones as they decompose, warming the water by an imperceptible amount. The amount of energy used and dissipated by marine life is too small to be singled out for special notice in figure 6.1. It forms a tiny fraction of the three "outgoing" arrows on the right.

Evaporative cooling entails the withdrawal of thermal energy from the sea-water and its conversion into latent heat in the water vapor that forms over the sea (see chapter 5). Occasionally energy travels the other way: water vapor in the air condenses to water at the surface of the sea, liberating latent heat that warms the water. But on average the net transfer of latent energy is from the sea to the air.

Likewise, heat can be conducted in either direction, depending on whether the sea at the surface is warmer or cooler than the air in contact with it. In other words, warmth—thermal energy—can be transferred downward from warm air to cool sea, or upward from warm sea to cool air, by conduction. In the latter case convection speeds up the transference, provided the temperature of the air falls at progressively greater heights above the surface; if it does not—if there is an inversion (the air temperature rising at greater heights)— convection can carry the warmed air only a short distance up. In any case the net transference of heat is upward, making conduction one of the ways—the least important way—the sea dissipates heat.4

The warmth of the sea is also transported horizontally, by currents; as with similar shifts of thermal energy in the atmosphere, such horizontal movements are called advection. Advection is not shown in figure 4.1 because directions and distances vary tremendously from place to place. All that can be said is that, over time, the advective heat transfers at any one location always cancel each other out.

How Sunlight Penetrates the Sea

As a whole, the ocean absorbs far more solar energy than does the land. This is scarcely surprising when you consider that the earth's surface is 71 percent ocean and 29 percent land. The ocean also absorbs more than the land per unit area, because of its lower albedo. Seawater is much less transparent than the atmosphere, however, and this limits the depth to which sunlight can penetrate.

Seawater also filters the sunlight shining through it.5 The water is more transparent to blue light than to other colors of the visible spectrum and almost opaque to invisible radiation, both infrared and ultraviolet. More than half the incident sunlight, including all of its infrared component, is absorbed in the topmost meter of the sea. By 10 m down, all colors but blue are gone. This is the reason the sea looks blue: the color comes from the light back-scattered by the molecules of seawater, and because all other colors are absorbed in the topmost layers, blue is the only one available for back-scattering from lower down. Reflection of the blue sky above makes a comparatively minor contribution to the blue of the sea on a sunny day.

Biologically productive water, containing mineral particles, biological pigments, and plankton, is much less transparent than pure water and usually looks greener.

The solar radiation shining into the sea warms it, making the surface water warmer than the water below it. This arrangement is very stable, in strong contrast with the state of affairs in the atmosphere, where sunshine heats the surface, which in turn heats the air in contact with it. This places warm air below cool air and leads to instability: convection currents of air—thermals— rise up from the hot ground, and in extreme cases, large air masses are overturned.

The way sunlit seawater heats up is strongly affected by its transparency.6 On a hot summer day the sun could, in theory, raise the temperature of the topmost 5 cm of clean seawater by 40°C. It doesn't happen because, during daylight, the sea is losing heat almost as fast as it gains it. It loses it by radiation, most rapidly from the very topmost skin of water, about 0.05 mm thick. The temperature within this skin is less at the top than at the bottom (in contrast to the state of affairs just below it), causing the thin skin of water to mix convectively and cool down; at the same time, turbulence in the air over the water induces matching turbulence, hence mixing, in the water itself. The rapid cooling prevents the theoretically possible (and incredible!) gain of 40°C that could happen were it not for the cooling; in practice, the temperature of the sea surface seldom changes by more than 1°C between day and night. The sea doesn't cool appreciably after sunset, whereas the atmosphere is cooling rapidly. It is this that makes the sea seem so warm when you swim at night; you are enjoying the stored energy of sunshine.

Movement in the Sea

The sea is never still. It moves in a variety of ways, from a number of different causes. Waves and tides are the movements most familiar to land dwellers, and they are the topic of chapters 7 and 8. In this chapter we consider ocean currents, the "winds" of the sea.

In the open ocean, far from any coastlines, the strongest ocean currents are those at the surface, caused by winds. Weaker currents at greater depths have a variety of causes. Last, some currents at the bottom of the ocean, known as hydrothermal currents, are caused by heat from the earth's interior; they are discussed in chapter 15.

Waves as well as surface currents are driven by wind energy. When the wind gets up over a calm sea, it drags along a thin skin of surface water, simultaneously starting a shallow current and creating tiny ripples. The ripples catch the wind and become small waves; this increases the wind drag, so that the waves continue growing in a feedback process. The higher the waves, the deeper the troughs between them, so the more deeply the wind can reach into the water and the thicker the current layer becomes. Despite appearances, the winds spend considerably more energy driving currents than they do raising waves.7

Although ocean currents resemble atmospheric winds in many ways, there are striking contrasts between them. Take the matter of speed: one writer has said that "in nature, air flows are normally about fifteen times as rapid as flows of water. Thirty meters per second is a hurricane of air; two meters per second is a torrent of water."8 At the slow end of the scale, the speeds at which flow is perceptible without instruments are in about the same ratio: the threshold of perceptibility is about 30 cm s-1 for an air current and about 2 cm s-1 for a water current.

The difference is due to the different viscosities of air and water.9 The viscosity of a material—be it a gas, a liquid, or a paste—is a measure of the force required to maintain a shearing motion of given speed in the material. To take a culinary example, envision an ice-cream sandwich, consisting of a flat slab of ice cream between two wafers. Suppose the bottom wafer is held firm on a horizontal surface; now you push the top wafer horizontally until it is sliding at a constant speed. For an ice-cream slab of given dimensions, the speed of the shearing motion depends only on the force you apply and the ice cream's viscosity: knowing the speed, the force, and the slab's dimensions, you can com pute the viscosity of the shearing movement in the ice cream. As you would expect, it is high when the ice cream is frozen hard and low when it is soft.

Let's return to the contrast between air and water. The viscosity of seawa-ter is about sixty times as great as that of air at sea level.10 No wonder their typical speeds of flow are so different. Their kinetic energies per unit volume are different too. For example, here are the speeds and energies per unit volume of three representative ocean currents:11

Current Speed (m s 1) KE (J m 3)

Florida Current (extremely fast)



Gulf Stream off Europe (slow)



Just strong enough to feel



And here, for comparison,

are the speeds and energies of representative winds:


Speed(ms 1)

KE (J m 3)

Violent storm (great damage)



Light breeze (leaves rustle)



Just strong enough to feel



As you can see, ocean currents have a higher kinetic energy per unit volume than comparable winds; this is because of the much greater density of water. The densities of seawater, and of air at sea level, in kilograms per cubic meter, are 1,024 and 1.2, respectively; that is, seawater is about 850 times as dense as air.13 Note, too, that the energy in ocean currents is far more concentrated. As we saw in chapters 4 and 5, wind speeds increase with height above the ground, so that air at all levels is in motion. In contrast, the speed of ocean currents decreases rapidly as you move downward from the surface; as we shall see later, nearly all the wind-driven movement is confined to a shallow surface layer, ranging from 100 m to 500 m thick.

Wind-Driven Currents and the Ekman Spirals

We have noted already that the strongest ocean currents are those at the surface, powered by the wind. The proof that the wind controls them is that they respond within a few hours to changes in wind direction.14

The speed of a wind-driven current is about 3 percent or less of the speed of the wind driving it. Much of the wind's energy is used up in the friction (strictly speaking, viscous drag) by which the wind drags the water along. Drag between wind and sea at the ocean surface and between layers of water

Figure 6.2. Map of how the winds above the ocean, and the surface current, respond to the Coriolis effect in the Northern Hemisphere: (a) atmospheric winds; (b) the wind and the current at the surface. See text for details.

Geostrophic wind a

Surface current b

Figure 6.2. Map of how the winds above the ocean, and the surface current, respond to the Coriolis effect in the Northern Hemisphere: (a) atmospheric winds; (b) the wind and the current at the surface. See text for details.

below the surface converts much of the mechanical energy of winds and currents to waste heat—entropy—an item not to be omitted from a complete energy budget.

The direction of a wind-driven current does not coincide with that of the wind driving it. This is another manifestation of the Coriolis effect described in chapter 5, which causes a surface current to be deflected through an angle of about 45° to the right of the wind. An angle of exactly 45°, which is what the simplest mathematical model of ocean circulation predicts, is not to be expected in the real world; the simple model assumes that the only factors determining a current's direction are the wind and the earth's rotation, whereas in reality a number of other factors influence the outcome. Assuming the simple model to be correct, figure 6.2 compares what happens in the lower atmosphere with what happens at the sea surface (we consider later what happens just below the surface). The figure applies to the Northern Hemisphere; as always happens with processes dependent on the earth's rotation, the pattern of events in the Southern Hemisphere is the mirror image of those in the Northern Hemisphere.

Figure 6.2a shows events in the atmosphere. The dotted arrow pointing north shows the direction the wind would blow given a nonrotating earth and a gradient of pressure decreasing to the north as it usually does. The solid

Figure 6.3. Ekman spirals at (a) the surface and (b) the bottom of the ocean. In both diagrams the open arrows low down point upward from the page. Note the different scales.

arrow pointing east is the geostrophic wind (see chapter 4), deflected 90° to the right by the Coriolis effect acting in the absence of drag. The dashed arrow pointing northeast shows the wind at the sea surface, where the Coriolis effect has been diminished by drag (see chapter 5); only half the deflection that produced the geostrophic wind remains, which makes it appear, as you descend from the heights, that the geostrophic wind has been deflected 45° to the left.

This is the wind that drives the current directly below it, as shown in figure 6.2b, where we see that the current at the surface (wavy arrow) has been deflected 45° to the right of the wind driving it. Behold, the surface current flows in the same direction as the geostrophic wind thousands of meters overhead, above the atmosphere's friction layer.

The next question is, How does the current flow at progressively greater depths below the water surface? The answer is given in figure 6.3a, which shows the directions and speeds of currents forming the Ekman spiral, so called in honor of the oceanographer who developed the first mathematical model of ocean circulation.15 The figure shows the water as a stack of horizontal layers, with the current in each layer moving more slowly than that in the layer above and somewhat to the right of it because of Coriolis deflection; in practice, of course, the layers are infinitesimally thin and blend into each other. The current peters out gradually, spiraling clockwise all the while. At the depth where its direction is directly opposite to the direction at the surface, known as the Ekman depth, the speed is only 4 percent of what it was at the surface.16 The water between the surface and the Ekman depth is the Ekman layer, which corresponds to the friction layer in the atmosphere.

The thickness of the Ekman layer depends on the wind speed and the latitude: the higher the latitude, the thinner the Ekman layer. For example, at latitudes 10°, 45°, and 80° (in either hemisphere), the Ekman depths would be 100 m, 50 m, and 45 m, respectively, given a fresh breeze of 10 m s-1, that is, a wind strong enough to raise moderate waves with many whitecaps.17 Doubling the wind speed to 20 m s-1 (a fresh gale) doubles the Ekman depths. Clearly, the Ekman layer is extremely thin compared with the friction layer of the atmosphere.

Now notice what seems at first sight a surprising fact: as you descend from high in the atmosphere, wind direction twists to the left;18 but as you descend from the sea surface into the depths, current direction twists to the right. Why the difference? Here is the answer. When the Coriolis effect twists the direction of a current of air or water, it twists it to the right of the direction in which it is being driven by the action of an external force, as explained in chapter 4. In the atmosphere, the external force is the atmospheric pressure gradient, exerting a force to the north in figure 6.2a (dotted arrow). In figure 6.2b the driving force is the wind, exerting a force to the northeast. Looked at in this light, it is clear that there is no difference; the two halves of figure 6.2 can be seen to correspond if you conceal the geostrophic wind arrow in 6.2a; then, in the two diagrams, you see arrows showing the initial causes and final outcomes of, respectively, a pressure gradient in the atmosphere and a wind driving the sea surface. In both cases the effect is directed 45° to the right of the cause.

Because of the constantly turning current direction as you descend below the sea surface, it follows that the bulk of the water is not moving in the same direction as the surface current; in fact the average direction of flow is at right angles to the wind direction, equivalently 45° to the right of the current at the surface. The average current is called Ekman transport; it is the flow that matters when we come to consider the transport of thermal energy by the ocean, though it is of no importance to sailors.

Other Ekman spirals form at the bottom of the ocean, where slow deepwa-ter currents are finally braked to a stop by the drag of the seafloor (fig. 6.3b). Because the currents are weak, the Ekman layers are thin; at 45° latitude and a current speed of 0.1 m s-1, the top of the Ekman layer is only 50 cm above the seafloor. The direction of a sea-bottom spiral matches the direction of the atmospheric Ekman spiral; that is, it twists to the left. This is as you would expect. In both cases, a geostrophic current of air (fig. 6.2a) or water (fig. 6.3b) is slowed by friction with a surface below it.

Returning to the Ekman layer at the top of the ocean, it is the layer in which virtually all wind-driven currents flow ("virtually" because, according to the model, wind-driven currents still have 4 percent of their strength at the Ekman depth).

The bulk of the energy imparted to the sea by the winds is used up in the surface Ekman layer. It is dissipated by viscous shearing as layers of water slide over each other. Viscosity exists in two forms: molecular viscosity and eddy viscosity.19 In molecular viscosity, individual molecules of water pass up and down between adjacent layers of water: then, whenever a fast molecule moves down among the slower molecules in the layer below it, it is slowed by collisions with the slow molecules and at the same time imparts some of its speed to them. Conversely, whenever a slow molecule moves up among the faster molecules in a layer above it, it is speeded up by collisions with the fast molecules, slowing them in the process. In this way the speed differences are evened out. In eddy viscosity the mechanism is the same, but the objects exchanging energy are big chunks of water instead of individual molecules. Eddy viscosity dissipates from 107 to 1011 times as much energy as molecular viscosity does. Eddy viscosity is therefore vastly more important in slowing ocean cur-rents.20 Realizing this, in 1902 Ekman succeeded in developing the first useful model of ocean circulation. Practically all the more recent models are refinements of Ekman's model made possible by high-speed computers.

Hills and Dales of the Ocean's Surface

Disregarding the small-scale ups and down of the waves, the surface of the ocean is not horizontal; in other words, sea level is not level, in spite of ap-pearances.21 A three-dimensional model of an ocean, with an enormously exaggerated vertical scale, would show a surface of smoothly sloping hills and valleys. The slopes are much too gentle to be directly observed from the surface, and their existence can only be inferred, usually from satellite observations. Their presence shows that gravitational potential energy (PE) is to be found at sea as well as on land, though in comparatively tiny amounts. Just as a rock at the top of a precipice has PE relative to the lowlands below, the water

Figure 6.4. The Pacific Ocean (a) under normal conditions and (b) during an El Niño event. Stippling shows warm water (about 30°C or more); arrows show currents.

atop a "sea hill" has PE relative to the lower surfaces surrounding it; the PE is converted to KE when water flows down the hill.

The tides are the most obvious cause of "hills" and "valleys" at sea. This follows directly from common sense, with no need for satellite data. The tides rise and fall because the whole body of ocean water encasing the earth, except where continents protrude, is not a spherical shell but a more or less football-shaped shell. The tides rise and fall because two bulges of water on opposite sides of the earth—the ends of the football—rotate around the earth. The bulges are low hills of water, each having a basal diameter the same as the earth's diameter. If you watch the tide rising at the beach, you are seeing one of these hills coming toward you; and if you begin your vigil at water's edge at low tide and don't get out of the way, the hill engulfs you. The evidence that the sea surface slopes could hardly be more convincing. (For more on tides, see chapter 8.)

Smaller hills form, superimposed on the tidal hills.A common cause is a fall in atmospheric pressure; if the pressure drops by 3 percent—equivalent to a drop from "set fair" to "change" on a household barometer—the sea level rises by about 30 cm.22 A rise or fall in sea level accompanying a fall or rise (respectively) in air pressure is known as the inverted barometer effect. The flooding that often accompanies severe hurricanes is a manifestation of the effect.

The sea also piles up where a current is halted by a barrier. For example, sea level in the tropical western Pacific is normally higher than in the tropical eastern Pacific because the westward-flowing South Equatorial Current piles up on reaching Indonesia and northern Australia. On the eastern side of the Pacific, cool water flows into equatorial latitudes from the south to replace that flowing west, causing a marked temperature gradient as well as a slope; the result is heaped-up warm water in the west and cool water at a lower level in the east; the difference in elevation across the width of the Pacific is about 50 cm. But every few years, air pressure rises over the western Pacific and falls over the eastern Pacific; the trade winds weaken, and the piled-up warm water in the west flows back into the eastern Pacific, raising surface temperatures by 8 or 9°C (fig. 6.4). This is an El Niño event; it is a rearrangement of thermal energy on a geographic scale, and it causes climatic havoc.

Currents also cause water to pile up even in the absence of land barriers. Indeed, wherever currents flow, the ocean surface slopes. The slopes are imperceptible: for instance, at 45° latitude, a current of 1 m s-1 causes a vertical difference in sea level of about 1 m in every 100 km, that is, a slope of 1 in 100,000.

The broadest "hills" in the ocean are enclosed in the huge "ring" currents known as gyres. As examples, figure 6.5 shows the gyres in the northern and

Figure 6.5. The Atlantic Ocean, showing its northern and southern subtropical gyres, NSG and SSG, and the currents circulating around them. (Other currents are not shown.) The dotted arrows show the prevailing winds: northeast and southeast trades on either side of the equator and "westerlies" at higher latitudes in both hemispheres.

Figure 6.6. An upwelling. The view is northward, into the prevailing wind (stippled arrow), along the west coast of California. The hollow arrows in the sea show Ekman transport (ET), powered by the wind, and an upwelling (UW) dragged up from deeper water. The slope of the sea surface is exaggerated.

Figure 6.6. An upwelling. The view is northward, into the prevailing wind (stippled arrow), along the west coast of California. The hollow arrows in the sea show Ekman transport (ET), powered by the wind, and an upwelling (UW) dragged up from deeper water. The slope of the sea surface is exaggerated.

southern Atlantic. The currents are driven by the winds and deflected by the landmasses. But that is only part of the explanation for them and doesn't account for their constancy: additional mechanisms, with feedback, are at work. Coriolis deflection continuously turns the currents, and it acts more strongly on warm water at the surface than on the cooler water below because the warm water is less dense. As a result, warm water piles up to form a hill within the gyre;23 in the North Atlantic, the surface at the summit of the hill is about 1.5 m higher than the surface at the periphery.

The water on the "hillslopes" has potential energy by virtue of altitude, small though it is. Water flows downslope under the pull of gravity and is deflected because of the Coriolis effect until it is flowing very gently downhill almost parallel with the contours—clockwise in the Northern Hemisphere, counterclockwise in the Southern Hemisphere. It behaves as the air does around an atmospheric anticyclone. The currents around the "hill" are almost geostrophic (controlled solely by gravity and the Coriolis effect) because the drag is very slight. They reinforce the wind-driven currents that were their initial cause.

Because the currents around gyres are maintained by two processes acting in concert, they are very constant. They do not respond nearly as quickly to changes in wind direction as do currents driven wholly by wind. Energy to keep the currents flowing is stored, as potential energy, in the topography of the water surface. It is released—converted to kinetic energy—gradually over many days. This makes the currents immune to short-lived wind changes.

Valleys as well as hills contribute to the relief of the sea surface. They appear where surface currents diverge from each other because the winds driving them diverge. Subsurface water flows upward to make good the loss.

Slopes also form where a current flows away from a coastline. This happens, in either hemisphere, wherever the prevailing wind blows toward the equator beside the west coast of a continent (see fig. 6.6). The wind drives a current and, as always, the Coriolis effect goes to work. The average flow is Ekman transport moving the bulk of the water at right angles to the wind di-rection—directly away from the land. The water surface develops a slope going uphill away from the land, and an ascending current of water from the depths flows upward to take the place of the water that Ekman transport has removed. This is an upwelling. The upwelling water is cooler than surface water, but not really cold: it rises gently, at about 10 m per day, from a shallow depth of 300 m at most.

Upwelling explains the surprising coolness of inshore waters off the west coasts of North and South America and southwestern Africa. Californians in particular are well acquainted with upwelling: along the shore at Cape Men-docino, the surface temperature of the water in August is about 7°C lower than the temperature 1,500 km out to sea at the same latitude.24

Density Currents in the Depths

Up to this point we have been concerned chiefly with the surface layer of the oceans, where wind-caused currents predominate. The layer is from 100 m to 500 m thick and is called the mixed layer because it is continually stirred by currents, including those that flow upward and downward. The mixed layer contains only about 2 percent of the whole ocean.

Below it, beyond the reach of the winds, is a somewhat thicker layer known as the pycnocline (from the Greek pyknos, thick, referring to the density of the water), in which the density increases rather abruptly. The increase in density is caused by a cooling of the water, an increase in its salinity, or both.25 Whichever it is, the change marks the level at which, going downward, wind-caused mixing stops and relative calm prevails. Below the change, in the deep zone, the density usually increases very gradually down to the bottom;26 the average temperature and salinity in the deep zone are 3.5°C and 34.7 parts per thousand.

In the deep zone, currents flow wherever there happen to be horizontal density differences. A density difference arises wherever there is a change in temperature or salinity; this creates a pressure gradient, down which water flows as a gentle current. Such currents are known as thermohaline currents (from the Greek therme, heat, and halos, salt).

Their energy is far less than that of wind-caused currents: their speed is only 1 or 2 km per day on average.27 But they are of great importance as transporters of thermal energy. At some locations, comparatively shallow thermohaline currents descend to become deep currents and then return to the surface; here and there they fork and subsequently rejoin. Taken all together, they form closed loops of enormous extent, which transport thermal energy from ocean to ocean and from the tropics toward the poles. These are thermohaline currents on the grand scale, functioning as energy conveyor belts.

The Global Energy Conveyor

The loop current that, with offshoots, spans the whole earth is shown schematically in figure 6.7. It has been dubbed the "global conveyor." The map in the figure combines the representations of several authors and cannot be correct in every detail.28 This should cause no surprise; the needed data are difficult to collect, and data points are often far apart. The cold, salty currents are at great depth, and the returning warm currents tend to spread out and become less definite as they rise toward the surface. Bear in mind that the conveyor is a three-dimensional structure: the cold currents are close to the ocean floor and ascend slowly to higher levels, warming as they rise. They also vary in salinity. That is, they are thermohaline currents, independent of the winds.

The deep, cold, salty current flowing from north to south for the length of the Atlantic Ocean is the start of the conveyor, insofar as it can be said to have a start. It is believed to come into existence in the following way: The Atlantic is a comparatively narrow ocean. Drying winds from nearby lands cause the

Figure 6.7. The "global conveyor" somewhat idealized. Deep, cold currents are heavily stippled; warmer currents, rising toward the surface, are lightly stippled. Atlantic water is saltier than Pacific water.

ocean to lose more water by evaporation than it gains from rainfall and from inflowing rivers. This makes North Atlantic water much saltier than water at the same latitude in the Pacific.29 At the same time as it loses heat by evaporation, the water is cooled by cold winds blowing from the Canadian Arctic. The upshot is that surface water in the vicinity of Iceland becomes steadily colder and saltier—and therefore denser—until it sinks, initiating the conveyor. It flows south and around the tip of Africa.

When it eventually reaches the South Pacific and turns north, the water lost through evaporation in the Atlantic is restored by excess rainfall in the Pacific: the salty water is diluted. Indeed, the conveyor carries salt as well as warmth: it evens out the salinity contrast between the two oceans.

The conveyor's role as a transporter of thermal energy is even more important. In the North Atlantic, before sinking, the conveyor "gives off a staggering amount of heat [to the atmosphere, which] accounts for the surprisingly mild winters of Western Europe."30 As figure 6.7 shows, in the Atlantic the global conveyor conveys heat from south to north for almost the whole length of the ocean, but in the Pacific the heat flow is poleward both north and south of the equator.

Ocean currents are indeed as important as the winds in carrying warmth from low latitudes to high. But it is difficult to compare the importance of ocean currents and atmospheric winds in achieving the redistribution of energy. According to one estimate, currents are more important than winds in the Northern Hemisphere south of latitude 25° N, whereas at higher latitudes the winds become more important.31 Depending on the season, however, the surface of the sea is sometimes cooler than the air above it and sometimes warmer, causing warmth to pass repeatedly from air to sea and back again. The atmosphere and the ocean act together in spreading warmth from the tropics to the polar regions.

Was this article helpful?

0 0
Getting Started With Solar

Getting Started With Solar

Do we really want the one thing that gives us its resources unconditionally to suffer even more than it is suffering now? Nature, is a part of our being from the earliest human days. We respect Nature and it gives us its bounty, but in the recent past greedy money hungry corporations have made us all so destructive, so wasteful.

Get My Free Ebook

Post a comment