Climate Change: The Science of Global Warming and our Energy Future

HOW CLIMATE WILL CHANGE in response to the buildup of greenhouse gases in the atmosphere will depend in large part on what happens in the world ocean. The reasons for this connection begin with the large amount of solar energy the ocean absorbs and holds, about 1,000 times more heat than the atmosphere holds. Consequently, the ocean is effective at both distributing that heat around the globe and moderating regional temperature (which is why summers in oceanic climates are not too hot or winters too cold). Moreover, the ocean also holds much more carbon than the atmosphere does. And because the two bodies readily exchange carbon dioxide (CO₂), the ocean exerts the primary control on the CO₂ content of the atmosphere and thus on climate, on timescales ranging from a century to a millennium and longer. For good reason, the ocean has been called the “caldron of climate.”¹

The ocean is immense, especially when considering its depths. Being mostly beyond view and unfamiliar, those depths are mysterious. Vast regions are more than 4,000 meters (13,000 feet) deep—as deep as the highest mountains are high. The ocean bottom is for the most part a broad, flat, featureless expanse known as the abyssal plains. The edges of the continents extend into the ocean as the continental shelves. In some places, these shelves reach out hundreds of kilometers from the shoreline. The entire North Sea, for example, is really part of the continent and is hardly anywhere more than 90 meters (300 feet) deep. Elsewhere, the shelf extends only a few tens of kilometers from shore—for example, along the west coast of South America. Then what is called the continental slope plunges steeply from the shelf edge to the deep ocean. The continental slope is a factor in guiding ocean circulation.

This chapter begins with the basic properties of water, which ultimately explain why the ocean exerts such an important control on the heat and carbon budgets of Earth’s climate system. It continues with an exploration of the ocean’s structure and circulation, both the surface currents, which are largely tied to the winds, and the circulation of the deep ocean, which can have a profound, long-term influence on climate. Finally, the chapter addresses the coupled interactions between the ocean and the atmosphere, using as an example the all important combination of the phenomena known as El Niño and Southern Oscillation, which affects weather patterns around the world.

It may seem odd to begin a discussion of the ocean with such an obvious and ubiquitous material as water (H₂O) itself, but water has some very special properties relevant to the ocean’s character. Imagine the H₂O molecule as an oxygen ion with two rabbit ears of hydrogen ions (figure 3.1). The oxygen ion has a negative electric charge, whereas the hydrogen ions have positive charges, making a water molecule “dipolar” (it has electrically negative and positive sides). (An ion is a charged particle that forms when a neutral atom or cluster of atoms, known as a molecule, either loses or gains one or more electrons to acquire either a positive [+] or a negative [–] charge.) This dipolarity causes individual water molecules to aggregate, with the positive side of one loosely attached to the negative side of another. This feature makes liquid water relatively stable. The uneven distribution of charge also enables water molecules to latch onto other ions, and for this reason water is an excellent solvent for many substances.

The stability of water means that the amount of energy needed to evaporate water and the amount of energy that must be removed to freeze it are among the highest of any substance. In other words, evaporation absorbs heat, and freezing releases it. Furthermore, the specific heat of water, which is the amount of energy required to raise a unit mass of water 1°C (1.8°F), is also among the highest of all materials. Water can thus absorb and release relatively large amounts of heat with very little change in temperature. Water’s high thermal capacity has an important consequence as far as climate is concerned. It results in the ocean being a huge reservoir for heat, holding, as noted earlier, about 1,000 times more heat than the atmosphere. So even though the ocean circulates much more slowly than the atmosphere, it moves a significant amount of heat around the planet.

The water molecule’s oxygen side has a small net negative charge, and its hydrogen side has a small net positive charge, allowing it to aggregate with other water molecules and other polar molecules. Because of these characteristics, liquid water is more stable than it would otherwise be and is an excellent solvent for many substances.

Water’s high specific heat has other consequences: the temperature of the ocean changes more slowly in response to heating or cooling than does the temperature of the atmosphere, and the total range and seasonal changes in ocean temperature are much less than those on land. To appreciate this, consider that the highest and lowest land temperatures, which have been recorded in the Libyan Desert and at the Russian Vostok research station in central Antarctica, respectively, are 58°C (136°F) and −88°C (−126°F), a range of 146°C (262°F). Ocean temperature, in contrast, varies from a maximum of about 36°C (97°F) in the Persian Gulf to −2°C (28°F) near the poles, a range of only 38°C (69°F). Annual sea-surface temperatures vary by no more than 2°C (3.6°F) in the tropics, 8°C (14.4°F) in the middle latitudes, and 4°C (7.2°F) in the polar regions, whereas seasonal temperature variations of more than 50°C (90°F) are not uncommon on continents. That is why the presence of a nearby ocean spares most coastal regions the frigid winters and sweltering summers that some continental interiors experience at similar latitudes. The ocean is a natural thermostat—it releases and takes up heat on timescales of decades to centuries, whereas the atmosphere does so in days to weeks. This difference occurs because of the fundamental character of water.

Seawater is only 96.5 percent water by weight. The rest of it is mainly sodium chloride (NaCl) and other dissolved salts (figure 3.2). Why is the ocean salty? The simple answer is that every year rivers dump into the ocean 2.5 billion metric tons of material, including dissolved positively charged ions of sodium (Na⁺), magnesium (Mg²⁺), and calcium (Ca²⁺) leached from rocks (chapter 4). In contrast, chlorine (Cl^–) and sulfate (SO₄^2–), the important negatively charged ions in seawater, have accumulated over geologic time and originated mostly from gases that escaped from Earth’s interior during volcanic eruption. Other materials in the ocean include dust from deserts and various pollutants from human activities.

But this description raises another question. Rivers have added far more dissolved matter to the ocean than the ocean actually contains; furthermore, the compositions of the matter dissolved in rivers and of the matter found in ocean water are different. For example, the ratio of Na⁺ to Ca²⁺ in seawater is about 26, but in river water the ratio is only 0.4. Why this difference, and why is the ocean not even more salty than it is? The answer is that the dissolved material continually precipitates and accumulates as sediment, and some compounds have a longer residence time in the ocean than others. Thus Na⁺ resides in the ocean much longer than Ca²⁺ because marine animals, such as phytoplankton, remove Ca²⁺ to make calcium carbonate (CaCO₃) shells. When the animals die, the carbonate shells accumulate on the ocean floor to make carbonate mud, which eventually becomes limestone. In contrast, Na⁺ is removed mainly by clay particles. The particles also settle to form sediment, but this process is much slower than carbonate removal. The composition of ocean water has remained remarkably constant over geologic time because the various dissolved components have attained a long-term balance as they cycle through the various parts of Earth.

The term salinity refers to the total salt content of water. The salinity of seawater is about 35 parts per thousand (ppt), essentially equivalent to 35 grams (1.2 ounces) of salt per liter of water. The addition of salt to water decreases the freezing point and increases density. Consequently, seawater freezes at −1.8°C (28.8°F) rather than at 0°C (32°F) and has a density of 1.026 grams per cubic centimeter, compared with 1.0 for pure water. (Pure water reaches its maximum density at 3.98°C [39.2°F].) Increasing temperature, however, decreases density. Density and thus salinity and temperature are among the fundamental factors that determine how the ocean circulates and affects climate. For example, in certain polar regions water sinks because it attains a relatively high density, and in the process it removes CO₂ from the atmosphere and stores it in the deep ocean.

Several processes can cause salinity to change. Evaporation removes pure water and thus increases the salinity of the remaining water; precipitation and influx of fresh river water decrease salinity by dilution; freezing removes and melting of ice adds freshwater to increase or decrease salinity. These processes combine to make ocean salinity vary from one place to another (figure 3.3). The most saline parts of the open ocean correspond to the belts of high atmospheric pressure on either side of the equator (the tropical highs, also marked by the global belts of desert [chapter 2]), where evaporation exceeds precipitation. Salinity is lower at the equator and, aside from the tropical highs, tends to decrease toward the poles because of greater precipitation and lower evaporation in these regions. Low salinities also exist locally near the mouths of large rivers, whereas very high salinities develop in semienclosed seas in arid regions—the Persian Gulf, the Red Sea, and the Mediterranean Sea being prime examples.

The most saline parts (red) of the open ocean exist where evaporation is highest and precipitation lowest—at the subtropical belts of high pressure on either side of the equator. Salinity decreases (blue) toward the poles because of greater precipitation and lower evaporation in these regions. Salinities are locally low near the mouths of large rivers and high in semienclosed seas in arid regions, such as the Mediterranean. Salinity is given as total salt content of water in parts per thousand. (After National Oceanic and Atmospheric Administration, National Oceanographic Data Center, http://www.nodc.noaa.gov/)

The average temperature of the ocean is 3.6°C (38°F), but if you swim in it, it is pretty obvious that the surface part of it is much warmer than that.² Indeed, because the density of water depends on both temperature and salinity, the ocean is thermally and compositionally stratified (figure 3.4).

The ocean has three major layers. A thin, warm surface layer, termed the mixed layer, extends to about 20 to 200 meters (66 to 660 feet) in depth (the average is 70 meters [230 feet]). Here temperature changes little as waves and convective overturning keep the waters stirred. Over years to decades, the ocean’s thermal capacity is essentially that of the mixed layer because over these timescales the mixed layer mixes little with the colder water below it.

Below the mixed layer is the thermocline, a zone in which temperature decreases and salinity increases rapidly with depth. The thermocline’s base is about 5°C (41°F) and generally extends to 500 to 900 meters (1,600 to 3,000 feet) depth, but this depth varies from location to location and season to season. In winter at certain high-latitude locations, the thermocline can reach all the way to the ocean bottom, as occurs, for example, in the Norwegian-Greenland Sea during particularly cold spells. Conversely, cold water exists near or at the surface around parts of Antarctica, in which case there is neither thermocline nor mixed layer.

Below the thermocline is the deep zone, where salinity and temperature vary only slightly with depth. The deep zone constitutes about 65 percent of ocean water and at its deepest and coldest is about 2°C (35.6°F) (figure 3.5).

Ocean surface currents are driven mainly by the wind, so like the wind they are affected by the Coriolis force and form distinctive patterns dominated by subtropical gyres, or semicircular current systems, on either side of the equator (figure 3.6). The generally easterly (out of the east or west-directed) trade winds drive the westward-flowing arms of the subtropical ocean gyres near the equator. This flow, in turn, generates eastward-flowing equatorial countercurrents both at and below the surface. Near the poles, easterly winds drive smaller gyres, but in the Southern Ocean strong westerly winds drag water around the entire globe in the Antarctic Circumpolar Current.

Of course, the positions of the continents also determine current patterns. In particular, the continents deflect the westward-flowing arms of the subtropical gyres, transforming them into strong, poleward western boundary currents flowing generally parallel to the coastlines. Important from the point of view of climate is the well-known Gulf Stream in the North Atlantic Ocean (figure 3.7), but other oceans also have their western boundary currents. In the North Pacific, the Kuroshio Current flows past Japan; in the South Atlantic, the Brazil Current flows along the coast of South America. The western South Pacific and Indian oceans have the analogous East Australian and Agulhas (Mozambique) currents, respectively. These currents are fast. For example, the Gulf Stream flows at a rate of nearly 2 meters (6.5 feet) per second, and it takes only about a month for warm water from the tropics to reach the midlatitudes.

A basic instrument in oceanographic research, the CTD device measures temperature, salinity (by measuring electrical conductivity of the seawater), and depth as it is lowered into the ocean, thus providing a continuous profile of these properties with depth. The vertical tubes are water bottles that may be closed at specific depths. When the device is returned to the ship, the water it has captured is typically analyzed for oxygen, CO₂, nutrients, and other substances. The photograph was taken on the Woods Hole Oceanographic Institution research vessel Atlantis during an expedition to the northeastern Pacific Ocean in 1997. (Photograph by the author)

The currents are driven mainly by the winds and, like the winds, are affected by Coriolis forces. The continents deflect the currents into strong, poleward western boundary currents, the most important of which from the point of view of climate is the Gulf Stream. (After Hartmann 1994, with permission)

The flow (reds and yellows) amounts to about 30 million cubic meters (39 million cubic yards) of water per second. The Gulf Stream and other boundary currents transport an enormous amount of heat to the high latitudes. (Liam Gumley, MODIS Atmosphere Team, University of Wisconsin–Madison Cooperative Institute for Meteorological Satellite Studies, http://visibleearth.nasa.gov/view_rec.php?id=1722)

The western boundary currents are important for several reasons. First, they carry a great deal of heat—especially the Gulf Stream,³ which not only is fast but has an enormous mean flow of about 30 million cubic meters (39 million cubic yards) of water per second. Mention should be made of the common misconception that the Gulf Stream is primarily responsible for northern Europe’s relatively warm climate. It is true that the Gulf Stream delivers a great deal of heat to the North Atlantic, but North Atlantic surface water retains a substantial amount of summer heat anyway, and winds out of the west and southwest transport warm surface water from lower latitudes toward Europe,⁴ so it is really the entire warm ocean surface that keeps western Europe relatively warm.

The western boundary currents are also important in that they remove CO₂ from the atmosphere. Cold water dissolves more CO₂ than does warm water. Therefore, the warm waters carried poleward by the western boundary currents initially contain relatively little CO₂, but as the waters cool, they take up CO₂ from the atmosphere.

The eastern boundary currents arise for the same reason that the western boundary currents do. They include, for example, the California and Peru currents, which flow along the west coasts of North and South America, and the Benguela Current, which flows along the west coast of southern Africa. All these currents flow from high latitudes, carry cold water toward the equator, and then turn westward away from the coasts.

It is fortunate for those of us who like to eat fish that ocean currents do not flow only in the direction of the wind. This phenomenon was first observed by the Norwegian explorer Fridtjof Nansen (1861–1930), who with his comrade F. H. Johansen in their ship the Fram spent the winter of 1893/1894 frozen in the Arctic pack ice. Their purposes were to study the Arctic Ocean currents and to get to the North Pole (they never reached it, despite mounting a mad dash with dogs and sledges when they realized the currents would not take them there). As the ice carried them along, Nansen noted that instead of drifting with the wind, the ice pack drifted at an angle 20 to 40 degrees to the right of the wind direction. Based on Nansen’s observation, the Swedish oceanographer Vagn Ekman (1874–1954) developed a model showing how, due to the Coriolis effect, ocean surface water should move 45 degrees to the right of the wind direction in the Northern Hemisphere and 45 degrees to the left of it in the Southern Hemisphere.⁵

What does all this have to do with fish, though? An important consequence of Ekman transport, as this phenomenon is known, is the upwelling of water along coastlines. Atmospheric circulation in subtropical regions produces winds parallel to the west coasts of South America and Africa. By Ekman transport, the winds carry the upper few tens of meters of the ocean surface layer away from shore, where it is replaced by colder water from depths of 100 to 200 meters (300 to 650 feet). The upwelling water is rich in nitrates and phosphates and thus supports vigorous growth of plankton and therefore of fish—hence, for example, the anchovy and sardine fisheries off Peru and Ecuador.

Ekman transport can also drive currents toward shore, in which case water piles up and then sinks. It also operates in the open ocean, producing, for example, the cold tongue of water extending from South America into the equatorial Pacific.

Water circulates through all the major oceans in what has been termed the global ocean conveyor system.⁶ More esoterically, this phenomenon is known as thermohaline circulation because differences in water density drive the circulation, and, as we have seen, water density depends on temperature and salinity. The ocean conveyor starts with downwelling in the North Atlantic Ocean and the Southern Ocean around Antarctica. This water flows to and mixes with Pacific Ocean water, which returns as a shallow current to replace the downwelling water (figure 3.8). The current exerts a stabilizing influence on global climate over hundreds to thousands of years, but changes in the current are paradoxically implicated in abrupt swings in climate, especially in the North Atlantic.

In this system, also known as thermohaline circulation, the current starts with the downwelling of cold, dense water in the North Atlantic and Southern oceans around Antarctica. This deep water flows to and wells up in the Pacific Ocean and then returns as a shallow current to replace the downwelling Atlantic Ocean water. This current exerts a major influence on global climate over hundreds to thousands of years. (After Intergovernmental Panel on Climate Change 2001: fig. 4.2)

The ocean conveyor works like this: warm, near-surface water in the Atlantic Ocean at about 35°N forms with a salinity of 34 to 36.5 ppt and flows northward at a depth of about 800 meters (2,600 feet). In the Norwegian-Greenland and Labrador seas during the winter, the northward-flowing water rises to the surface as surface waters are displaced by winds. Upon reaching the surface, the new water, which is initially at a relatively warm 10°C (50°F), cools rapidly to about 2°C (35.6°F) as it loses heat to the atmosphere. The cooling increases the water’s density, so it sinks to become what is known as North Atlantic Deep Water (NADW). Now at a temperature of 2 to 4°C (35.6 to 39.2°F) and with a salinity of 34.8 to 35.1 ppt, NADW forms a deep, dense current that flows south, through the South Atlantic, and all the way to the Southern Ocean.

In the vicinity of about 60°S, NADW flows over Antarctic Bottom Water (AABW). The AABW originates mainly by downwelling in the Weddell Sea, where the densest water in the ocean forms due to the cold and increased salinity brought about as winter sea ice forms. It spreads as tongues into the Pacific, Indian, and Atlantic oceans, but in the Atlantic it mixes with and becomes entrained in the much stronger southward flow of NADW, which eventually spreads out through the Indian Ocean and into the southern Pacific Ocean.

The cold bottom water may make it as far as 30°N latitude in the Pacific Ocean, but eventually it mixes with the water above it. Pacific Ocean water returns to the Atlantic Ocean as a warmer, less salty current at shallow to intermediate depths through the Indian Ocean and around the southern tip of Africa. The amount of water flowing in the global ocean conveyor system is about 15 times that flowing in all the world’s rivers, and the deep water mixes through the world ocean in a period of about 1,000 years.

The northward flow of water in the North Atlantic, also known as meridional overturning circulation, represents an enormous transport of heat—in fact, about one-quarter of the total global meridional (that is, directed north to south) heat transport. This explains why winters in the North Atlantic, despite the high latitude, are relatively mild, and it illustrates how the ocean can have a significant moderating effect on regional climate. Geological records of past climate indicate that on a number of occasions the thermohaline circulation has slowed or stopped and then restarted abruptly (that is, within a few years), implicating it in some of the rapid and dramatic shifts in climate experienced in the past by Greenland and northern Europe (chapter 6).

The global ocean conveyor system is intimately coupled to water transport through the atmosphere. Over the whole of the North Atlantic, there is more evaporation than precipitation (table 3.1), so the air leaving the Atlantic is more moist than the air entering it, which makes Atlantic water saltier than North Pacific water (by about 5 percent) (see figure 3.3). This extra saltiness is why deep water forms in the North Atlantic rather than in the North Pacific. The Pacific, in contrast, experiences slightly more precipitation than evaporation.⁷ The global ocean conveyor acts to off set this imbalance by redistributing the salt throughout the world ocean.

The Mediterranean Sea is also implicated in this affair. Mediterranean water is relatively saline because of a high evaporation rate. (The salinity is 37 to 38 ppt, which is why the lazy swimmer floats so easily in it.) Freshwater flows into the Mediterranean from several major rivers (most important the Nile), and less-saline Atlantic Ocean surface water flows in through the Strait of Gibraltar. But beneath the incoming surface flow, the more saline Mediterranean water flows out through the strait in the reverse direction. In the Atlantic, this saline water sinks, mixes, and spreads for thousands of kilometers. The water from the Mediterranean increases the salinity of the North Atlantic by about 6 percent over what it would otherwise be.

As water flows around the globe, land experiences more precipitation than evaporation, but over the ocean the opposite is true (see table 3.1). The different continents receive significantly different amounts of water as precipitation and lose different proportions of water by runoff (the water that flows back to the ocean via rivers) and through the combined processes of evaporation and plant transpiration, termed evapotranspiration. The flow of water through the ocean, atmosphere, and land is known as the hydrologic cycle. The amount of water that moves through the hydrologic cycle every year is equivalent to about a 1-meter-deep (3-foot) layer of water over the entire Earth.

Every several years, a dramatic and rapid warming occurs in the waters off the coast of Peru and Ecuador. Within a month, water temperature increases by 2 to 4°C (3.6 to 7.2°F). Anchovies disappear; sardines move south; and birds, fur seals, and other animals whose livelihoods depend on the fish die. Heavy rains inundate the normally arid coastal regions of northern Peru and Ecuador and cause flooding, but droughts strike the Andes of southern Peru and more distant regions such as northeastern Brazil. The warming of the Pacific usually takes place around Christmas and persists into May or June. Local fisherman refer to it as El Niño, Spanish for “Christ Child.”

El Niño is an ocean phenomenon, but it is intimately connected to the atmospheric phenomenon known as the Southern Oscillation, which is a swing in barometric pressure (the weight of air around us) between the eastern and western parts of the equatorial Pacific Ocean. Because the two are coupled, they are now generally referred to together as El Niño–Southern Oscillation (ENSO).

As we have come to learn, ENSO is hardly limited to the Pacific Ocean and surrounding regions. Rather, it affects climate around the globe. It commonly creates severe drought in Australia, Indonesia, and southern Africa;⁸ weakens the Asian monsoons; and brings mild but stormier winter weather to North America. It even affects the North Atlantic, where hurricanes are markedly less common and less intense in El Niño years than in “normal” years. The effects on ecosystems and agriculture can be dramatic (figure 3.9). For example, the unusually strong 1997/1998 event resulted in the death of 16 percent of the world’s tropical corals.⁹

El Niño is the consequence of a change of atmosphere and ocean circulation across the entire equatorial Pacific and is a dramatic example of how the two interact. It occurs at irregular intervals of 2 to 7 years, appears on average about every 4 years, and has been a feature of the climate system for at least the past 130,000 years.¹⁰ Next to the seasons, ENSO is the most important recurrent change in climate.

The maize yield in Zimbabwe is a close reflection of rainfall. Sea-surface temperature is the average for February and March, the height of the growing season in Zimbabwe, for the region in the equatorial Pacific bounded by 90 to 150°W longitudes and 5°N to 5°S latitudes. The correlation demonstrates the effect that El Niño–Southern Oscillation has on distant parts of the world. The maize yield is in kilograms per hectare relative to national average. (Adapted by permission from Macmillan Publishers Ltd: Nature, Cane, Eshel, and Buckland 1994, copyright 1994)

How does ENSO work? In “normal” years, sometimes referred to as La Niña (the Girl) years, especially when the winds are strong, the easterly trade winds of the Northern and Southern Hemispheres converge along the equator to blow west across the equatorial Pacific Ocean, driving the westward-flowing equatorial current (figure 3.10). The easterly winds cause Ekman transport away from the equator and upwelling of cold water in a narrow band along it, while coastal winds cause upwelling along the west coast of South America. A strong, eastward-flowing, equatorial “undercurrent” provides the water to replace these upwellings. The westward-flowing surface water piles up in the western Pacific, causing sea level to be about 60 to 70 centimeters (23 to 27 inches) higher there than in the eastern Pacific. The upwelling brings the thermocline essentially to the surface in the east, and in the west the pileup of surface water causes the thermocline to be much deeper, typically about 150 meters (490 feet) below the surface.

In the meantime, ocean and land surfaces of the western Pacific heat the atmosphere. The warm, moist air rises to form strong thunderstorms in what is known as the Indo-Australian Convergence Zone (IACZ). The convergence occurs because the rising air, which causes the barometric pressure of the IACZ to be low, is replaced by easterly winds from the Pacific and westerly winds from Australasia (figure 3.11). The rising air feeds a returning, eastward-directed flow of air near the top of the troposphere that eventually sinks in the eastern Pacific, creating a zone of high barometric pressure. This equatorial convective cycling of the atmosphere has become known as Walker circulation, for British physicist Sir Gilbert Walker (1868–1958), who studied records of the Indian monsoon while stationed in India and recognized the relationship between the strength of the monsoon and the changing wind patterns in the Pacific Ocean.¹¹

In normal (La Niña) years (a), the winds blow west across the equatorial Pacific Ocean, causing upwelling of cold water as surface waters are continuously displaced through Ekman transport. In El Niño years (b), the winds weaken, and warm western Pacific water flows back toward the east. El Niño and La Niña events affect the jet streams and Walker circulation, thereby changing weather patterns in distant parts of the globe. Most prominently, El Niño years bring droughts to parts of Southeast Asia, South America, and Africa (dark brown). (Krishna Ramanujan, NASA Goddard Space Flight Center, http://www.nasa.gov/vision/earth/lookingatearth/elnino_split.html)

It is worth emphasizing the strong, positive feedback between the ocean and the atmosphere in this coupled circulation pattern:¹² the strong easterly surface winds maintain a warm western Pacific Ocean; the warm waters heat the overlying air and cause it to rise; the rising air results in low barometric pressure in the west and high pressure in the east; the consequent pressure gradient causes strong winds to blow west across the eastern equatorial Pacific; and the strong winds keep the warm water in the western Pacific. The ENSO phenomenon represents one of several coupled atmosphere–ocean circulation patterns that influence climate on timescales of years to several decades.

In an El Niño year, this coupled ocean–atmosphere circulation pattern of the equatorial Pacific collapses (see figure 3.10). The easterly trade winds of the eastern Pacific weaken, warm western Pacific water flows back toward the east, and sea level flattens out—in essence, the water that typically piles up in the west sloshes back in an enormous (but imperceptible to us) “wave.” Upwelling ceases along the coast of South America and along the equator, the thermocline deepens to tens of meters in the east, and the equatorial undercurrent stops. The warmer water in the east now heats the air over the eastern Pacific, causing the air to rise and bringing the heavy rains to coastal South America.

In the meantime, in the western Pacific the IACZ becomes less localized and shifts eastward. The regular Walker circulation of normal years breaks down, and descending dry air commonly falls over Indonesia and Australia to create drought conditions there. This change in Walker circulation is also the mechanism through which ENSO influences the Asian monsoon. The shift to El Niño also causes a shift in the trade winds (see figure 3.10), changing weather patterns in other, distant parts of the globe.

The equatorial Pacific oscillates between its El Niño and normal states probably because the depth of the thermocline in the east, which influences sea-surface temperature for a given wind velocity, does not change exactly in phase with changes in the winds.

The historical record of the Southern Oscillation Index reveals several important features. First, although ENSO events occurred on average about every four years from 1935 to 1995 and in that sense are quasi-periodic, their occurrence is highly irregular. For example, from 1978 to 1986 there was only one ENSO event, but from 1990 to 1995 there was either one prolonged event or several that followed each other in close succession. Second, some ENSO events are intense, while others are not. The 1997/1998 ENSO was perhaps the fiercest in the past 1,000 years, and the 1982 event was also unusually intense. Others, however, are barely noticeable. Third, ENSO states alternate with what we have been calling “normal states,” or La Niña events. In fact, La Niña and El Niño events can simply be viewed as two alternating states of ocean–wind circulation patterns, with neither one nor the other being abnormal.

The SOI is the deviation from the mean difference in atmospheric pressure between Tahiti and Darwin, Australia. Sea-surface temperature is the monthly difference from the mean in the region bounded by 5°N to 5°S latitudes and 120 to 170°W longitudes. Note the 1982/1983 and 1997/1998 events, which are thought to have been the strongest of the twentieth century. (After McPhaden, Zebiak, and Glantz 2006, with permission)

Needless to say, what ultimately causes ENSO has engendered and continues to engender debate; it also raises the question of what will happen to ENSO as climate warms. This question is fairly important because sea-surface temperature in the eastern tropical Pacific Ocean appears to affect climate, especially rainfall, in parts of western North America (chapter 7). A clear picture of how warming will affect ENSO or how ENSO will affect climate in a warmer world has yet to emerge.

This chapter has sought to show how the ocean is an integral part of the climate system for two primary reasons. First, the ocean stores far more heat than the atmosphere, so the two in concert move heat around the planet, primarily from the tropics to the poles. Second, the ocean also holds far more carbon than the atmosphere. More important, because the two exchange CO₂, the ocean also affects global climate by controlling the CO₂ content in the atmosphere over periods of a century and longer. In essence, the ocean acts as a global thermostat by being a large reservoir for both heat and carbon.

The tight physical and chemical coupling of the ocean and the atmosphere should now be evident. The present climate system displays this coupling by a variety of features. We encountered two features, the monsoons and North Atlantic Oscillation (NAO), in the previous chapter, but another one stands out, ENSO, which has a global reach and, next to the seasons, represents the most important recurrent change in climate. In succeeding chapters, we shall discover some of the roles these ocean–atmosphere interactions play in determining how global warming will affect conditions on the continents.