The Global Climate System
This chapter provides a basic overview of the global climate system. Global climate emerges from the complex interaction over multiple time scales of atmospheric, oceanic, and terrestrial subsystems. These interactions produce characteristic patterns of oceanic and atmospheric circulation, such as the El Niño Southern Oscillation, described in this chapter. Both internal variability and external forcings drive short-term variations and longer-term changes in climate, and these properties can render local climates hard to predict with precision.
What we call the Earth’s climate system consists of several subsystems. These interact with each other on very different timescales: the atmosphere over several thousands of kilometers can change substantially on daily and subdaily scales; the ocean currents vary over timescales of months to millennia; and the huge ice sheets change significantly on millennial timescales. Over even longer periods, other parts of the Earth’s system also come into play, such as plate tectonics, which modify the Earth’s surface by generating new ocean basins and mountain ranges and by moving the geographical position of continents. This characteristic of multiple systems and timescales renders the climate system hard to predict because myriad different physical processes have to be included to provide any realistic description of the whole.
The subsystems of the climate system—atmosphere, ocean, land ice, land vegetation cover, and so on—all with their variations on different timescales, interact through the exchange of energy and matter. In particular, greenhouse gases such as water vapor, carbon dioxide, and methane are constantly being exchanged; and when set free in the atmosphere, they significantly influence the balance between absorbed and emitted energy at the Earth’s surface. In this regard, water vapor, liquid water, and ice in the atmosphere deserve special consideration since they lead to the formation of several types of clouds each with different properties regarding the reflection and absorption of radiation.1
The interplay among these climate subsystems is strongly non-linear, so that some perturbations can be rapidly amplified once they arise. The Earth’s climate is an open system, absorbing shortwave radiation from the sun, which is then distributed among its subsystems until it is finally radiated back to space in the form of longwave (thermal) radiation. These non-linear interactions and the continuous flow of energy result in internal climate variability on all timescales. This variability would occur even if the orbit of the Earth and the output of the sun were constant, providing exactly the same external source of energy to the Earth’s climate: each year, each decade, each century would be different, each being but one sample of that probabilistic distribution that we call “climate.”
Most of the incoming solar energy is transformed at the surface, with some smaller portions absorbed in the troposphere (the lowest level of the atmosphere) and in the stratosphere (just above the troposphere). Therefore the lower portion of the atmosphere is a system that is mainly heated from below (heat radiating up from the land or sea surface), whereas the ocean is mainly heated from above (incoming solar radiation). Warm air is lighter than cold air and warm seawater is lighter than cold seawater. Since warmer air usually underlies cold air in the atmosphere, but warmer surface ocean water rests on top of colder subsurface water, we tend to find unstable and turbulent atmospheric dynamics, but generally stratified and stable oceans, especially in tropical and subtropical regions with high upper-ocean temperatures.
Moreover, the tilt of Earth’s axis (currently 23.5°) means that the zone of maximum solar insolation shifts from the northern tropics (in the Northern Hemisphere summer) to the southern tropics (in the southern summer). This alternation generates the annual cycle of thermal (hot and cold) and hydrological (wet and dry) seasons over most of the globe. In general, lower latitudes typically show hydrological seasons, whereas mid to high latitudes are characterized by a more or less pronounced seasonality in temperatures.
The poleward transport of heat by the atmosphere is framed by three circulation cells.3 The first is the Hadley Cell. Over low-latitude tropical areas, warm air rises. Once it reaches the upper troposphere (around 16 km above sea level) it is deflected towards the poles. As it moves towards the mid latitudes, the air descends into lower tropospheric levels creating large subtropical high-pressure cells. From these zones of high pressure, air flows back towards the equatorial regions in the form of more or less constant southeasterly trade winds, which blow into the low-pressure Inter-Tropical Convergence Zone. (Note that winds are named after the direction from which they flow, so an “easterly” blows from east to west.)
Second, at its poleward branches the Hadley Cell interferes with the Ferrel Cell. This cell, located mainly over the mid latitudes, is characterized by prevailing westerly winds. These come mainly from the deflection of the upper tropospheric air particles towards the east in the presence of the Coriolis force—that is, because the (west to east) rotation velocity of the Earth’s surface decreases from the equator to the poles. In addition, atmospheric turbulence causes the familiar transient low- and high-pressure systems of the mid latitudes. The Ferrel Cell accounts for a considerable poleward heat transport. Third, over high latitudes the air cools further and descends, forming large high-pressure cells over the polar regions. This movement creates the Polar Cell, with prevailing easterly winds.
The oceanic part of the energy transport is more strongly determined by the shape of the ocean basins. One important mechanism is the narrow western boundary currents flowing along the eastern side of major continents at mid latitudes, such as in North America (the Gulf Stream) and eastern Eurasia (the Kuroshio). These currents result from the interplay of three factors: the wind force provided by the semipermanent subtropical high air-pressure cells, the rotation of the Earth, and the generally longitudinal orientation of the coastlines. These narrow currents transport warm tropical waters polewards. The waters then generally flow back towards the equator along the eastern side of the ocean basin, forming much broader current systems, such as the Canary Current.
The Atlantic Ocean deviates from the Pacific Ocean in one important respect: in the North and South Atlantic at high latitudes, the surface waters are colder and more saline, and therefore denser. This density leads to “deep water formation”: deep convection that transports high-latitude cold water masses from the surface to the ocean interior, leaving them to be replaced by warmer water masses from lower latitudes. This poleward flow at the surface, known as thermohaline circulation, not only is another driver of poleward heat transport but also represents an important way in which warm surface waters and cold deep waters are mixed in the oceans, which are generally stratified—that is, layered between waters of different temperature.4 In this way, heat stored in the upper oceanic layers can penetrate down into the deep ocean, a mechanism that is important for controlling and mediating climatic changes on millennial timescales.
The geographical arrangement of the continents also results in particular regional climates in specific bands of latitude. One example is the Indian monsoon system, largely a result of the Himalayas and the Tibetan plateau being located close to the tropical Indian Ocean. A monsoonal climate is defined by a seasonal change in prevailing wind direction of at least 120°. With some simplification, the monsoon can be thought of as a sort of land–sea breeze but on a continental and seasonal scale. During winter, the Tibetan plateau cools down, giving rise to descending air masses and hence producing a pronounced high-pressure system and easterly winds. As the winds flow from continental areas, they carry little moisture, and precipitation is low (with the exception of the areas facing towards the Bay of Bengal). The summer monsoon, on the other hand, is driven by a strong low-pressure system developing over the Asian land masses owing to the higher heating rates over land during the (northern) summer season. This results in very humid southwesterly winds flowing from the Indian Ocean across the Indian subcontinent, bringing heavy seasonal rains and orographic amplified precipitation (i.e., precipitation enhanced by the rising of moist air as it passes over mountains) along the coastal ranges of the Ghats. Similar monsoon systems can be found in other parts of the tropics, including Africa, Southeast Asia, and North America.
Mean climate, as described above, represents only an average picture, not what is actually observed. At any particular point in time, we find configurations of the atmosphere, ocean, and cryosphere that are constantly varying within certain ranges around the mean climate state. In a stable climate, this variability is the result of numerous interactions within each subsystem and among the climate subsystems.
A paramount example of this internal variability is the El Niño-Southern Oscillation phenomenon.5 Usually, the easterly trade winds in the Tropical Pacific drive warm surface waters towards the west, triggering an upwelling of colder subsurface waters off Peru. This phenomenon maintains a temperature and surface pressure gradient across the whole Tropical Pacific, which in turn reinforces the trade winds. That is, the colder waters and higher air pressure in the Eastern Tropical Pacific and the warmer waters and lower air pressure in the Western Tropical Pacific help sustain the usual east-to-west winds. If for any reason the trade winds slacken, the temperature and pressure gradient also weaken, thus further weakening the trade winds. For a few months, about every five years or so, the whole Tropical Pacific shifts to this different “state,” called “El Niño,” when trade winds slacken and the Eastern Tropical Pacific becomes unusually warm. El Niños change surface temperatures, ocean vertical mixing, and surface heat fluxes so strongly that they may affect the atmosphere not only in the Tropical Pacific but also globally, via so-called “teleconnections.” Strong El Niño years are therefore associated with climatic effects as diverse as heavy rainfall in Peru and droughts in East Africa, India, and Australia (see Chap. 34).
The term “climate change” (as opposed to “climate variability”) denotes a modification in the statistics of the weather in the atmosphere—and, expanding the meaning of the concept of “weather,” also of the ocean and other subsystems. These changes can be brought about by various “forcings.” The term “forcing” denotes a driving factor that is considered to be external to the climate system. It may be embedded in the Earth’s system, as in the case of volcanoes, or be truly extraterrestrial, as in the case of the sun. Examples of external forcings include shifts in the configuration of the continents by plate tectonics (on geological timescales), variations in the output of the sun, volcanic eruptions, and anthropogenic emissions of greenhouse gases, such as carbon dioxide and methane.
All of these forcings at least temporarily disturb the balance of energy that is absorbed and released by the Earth. For example, continental masses at high latitudes allow the formation of permanent ice sheets. These increase the albedo (reflectivity) of the Earth’s surface, and a higher albedo means that more solar radiation is reflected back to space before it even enters the energy cycle of the climate system. Another example is the increase in atmospheric greenhouse gases. These gases hinder the release of longwave radiation from the Earth’s surface back to space, so that more energy becomes trapped within the climate system.
The climate system will adjust to such perturbations until a new energy balance is reached. In the first example, the surface temperatures will tend to cool, thereby emitting less longwave radiation to space and reducing energy losses. In the second example—the situation which we are currently in (see Chap. 26)—surface temperatures will tend to increase, thus radiating more thermal radiation upwards, compensating for the “trapping” effect of atmospheric greenhouse gases. These readjustments are accompanied by changes in atmospheric and oceanic circulation, cloud cover, atmospheric water vapor, and many other factors that in turn also affect surface temperatures.6 The theoretical term “climate sensitivity” summarizes all of these complex processes in a single number, which states the amount of surface warming that is required to achieve a new state of energy balance.
- Holton, J.R., and R. Dmowska. El Niño, La Niña, and the Southern Oscillation. Edited by S.G. Philander. San Diego: Academic Press, 1990.Google Scholar