/kluy"mit/, n.
1. the composite or generally prevailing weather conditions of a region, as temperature, air pressure, humidity, precipitation, sunshine, cloudiness, and winds, throughout the year, averaged over a series of years.
2. a region or area characterized by a given climate: to move to a warm climate.
3. the prevailing attitudes, standards, or environmental conditions of a group, period, or place: a climate of political unrest.
[1350-1400 for earlier senses; 1595-1605 for def. 2; ME climat < L climat- (s. of clima) < Gk klimat-, s. of klíma slope, equiv. to kli- (akin to klínein to slope, lean) + -ma n. suffix]
Syn. 3. mood, atmosphere, spirit, tone, temper.

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Condition of the atmosphere at a particular location over a long period of time (from one month to many millions of years, but generally 30 years).

Climate is the sum of atmospheric elements (and their variations): solar radiation, temperature, humidity, clouds and precipitation (type, frequency, and amount), atmospheric pressure, and wind (speed and direction). To the nonspecialist, climate means expected or habitual weather at a particular place and time of year. To the specialist, climate also denotes the degree of variability of weather, and it includes not only the atmosphere but also the hydrosphere, lithosphere, biosphere, and such extraterrestrial factors as the sun. See also urban climate.

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 conditions of the atmosphere at a particular location over a long period of time; it is the long-term summation of the atmospheric elements (and their variations) that, over short time periods, constitute weather. These elements are solar radiation, temperature, humidity, precipitation (type, frequency, and amount), atmospheric pressure, and wind (speed and direction).

      From the ancient Greek origins of the word (klíma, “an inclination or slope”—e.g., of the Sun's (Sun) rays; a latitude (latitude and longitude) zone of the Earth; a clime) and from its earliest usage in English, climate has been understood to mean the atmospheric conditions that prevail in a given region or zone. In the older form, clime, it was sometimes taken to include all aspects of the environment, including the natural vegetation. The best modern definitions of climate regard it as constituting the total experience of weather and atmospheric behaviour over a number of years in a given region. Climate is not just the “average weather” (an obsolete, and always inadequate, definition). It should include not only the average values of the climatic elements that prevail at different times but also their extreme ranges and variability and the frequency of various occurrences. Just as one year differs from another, decades and centuries are found to differ from one another by a smaller, but sometimes significant, amount. Climate is therefore time-dependent, and climatic values or indexes should not be quoted without specifying what years they refer to.

      This article treats the factors that produce weather and climate and the complex processes that cause variations in both. Other major points of coverage include global climatic types and microclimates. The article also considers both the impact of climate on human life and the effects of human activities on climate. For details concerning the disciplines of meteorology and climatology, see climatic variation and change (climate change). See also the article atmosphere (atmosphere, evolution of) for further information about the properties and behaviour of the atmospheric system. Relevant data on the influence of the oceans and of atmospheric moisture on climate can be found in hydrosphere.

solar radiation and temperature
       air temperatures have their origin in the absorption of radiant energy from the Sun. They are subject to many influences, including those of the atmosphere, ocean, and land, and are modified by them. As variation of solar radiation is the single most important factor affecting climate, it is considered here first.

Solar radiation
Distribution of radiant energy from the Sun
       nuclear fusion deep within the Sun releases a tremendous amount of energy that is slowly transferred to the solar surface, from which it is radiated into space. The planets intercept minute fractions of this energy, the amount depending on their size and distance from the Sun. A 1-square-metre (11-square-foot) area perpendicular (90°) to the rays of the Sun at the top of Earth's atmosphere, for example, receives about 1,365 watts of solar power. (This amount is comparable to the power consumption of a typical electric heater.) Because of the slight ellipticity of Earth's orbit around the Sun, the amount of solar energy intercepted by Earth steadily rises and falls by ±3.4 percent throughout the year, peaking on January 3, when Earth is closest to the Sun. Although about 31 percent of this energy is not used as it is scattered back to space, the remaining amount is sufficient to power the movement of atmospheric winds and oceanic currents and to sustain nearly all biospheric activity.

 Most surfaces are not perpendicular to the Sun, and the energy they receive depends on their solar elevation angle. (The maximum solar elevation is 90° for the overhead Sun.) This angle changes systematically with latitude (latitude and longitude), the time of year, and the time of day. The noontime elevation angle reaches a maximum at all latitudes north of the Tropic of Cancer (23.5° N) around June 22 and a minimum around December 22. South of the Tropic of Capricorn (23.5° S), the opposite holds true, and between the two tropics, the maximum elevation angle (90°) occurs twice a year. When the Sun has a lower elevation angle, the solar energy is less intense because it is spread out over a larger area. Variation of solar elevation is thus one of the main factors that accounts for the dependence of climatic regime on latitude. The other main factor is the length of daylight. For latitudes poleward of 66.5° N and S, the length of day ranges from zero (winter solstice) to 24 hours (summer solstice), whereas the Equator has a constant 12-hour day throughout the year. The seasonal range of temperature consequently decreases from high latitudes to the tropics, where it becomes less than the diurnal range of temperature.

Effects of the atmosphere
 Of the radiant (radiation) energy reaching the top of the atmosphere, 46 percent is absorbed (absorption) by Earth's surface on average, but this value varies significantly from place to place, depending on cloudiness, surface type, and elevation. If there is persistent cloud cover, as exists in some equatorial regions, much of the incident solar radiation is scattered back to space, and very little is absorbed by Earth's surface. Water surfaces have low reflectivity (4–10 percent), except in low solar elevations, and are the most efficient absorbers. Snow surfaces, on the other hand, have high reflectivity (40–80 percent) and so are the poorest absorbers. High-altitude desert regions consistently absorb higher-than-average amounts of solar radiation because of the reduced effect of the atmosphere above them.

      An additional 23 percent or so of the incident solar radiation is absorbed on average in the atmosphere, especially by water vapour and clouds at lower altitudes and by ozone (O3) in the stratosphere. Absorption of solar radiation by ozone shields the terrestrial surface from harmful ultraviolet light and warms the stratosphere, producing maximum temperatures of −15 to 10 °C (5 to 50 °F) at an altitude of 50 km (30 miles). Most atmospheric absorption takes place at ultraviolet (ultraviolet radiation) and infrared (infrared radiation) wavelengths (wavelength), so more than 90 percent of the visible portion of the solar spectrum, with wavelengths between 0.4 and 0.7 μm (0.00002 to 0.00003 inch), reaches the surface on a cloud-free day. Visible light, however, is scattered in varying degrees by cloud droplets, air molecules, and dust particles. Blue skies and red sunsets are in effect attributable to the preferential scattering of short (blue) wavelengths by air molecules and small dust particles. Cloud droplets scatter visible wavelengths impartially (hence, clouds usually appear white) but very efficiently, so the reflectivity of clouds to solar radiation is typically about 50 percent and may be as high as 80 percent for thick clouds.

      The constant gain of solar energy by Earth's surface is systematically returned to space in the form of thermally emitted radiation (thermal radiation) in the infrared portion of the spectrum. The emitted wavelengths are mainly between 5 and 100 μm (0.0002 and 0.004 inch), and they interact differently with the atmosphere compared with the shorter wavelengths of solar radiation. Very little of the radiation emitted by Earth's surface passes directly through the atmosphere. Most of it is absorbed by clouds, carbon dioxide, and water vapour and is then reemitted in all directions. The atmosphere thus acts as a radiative blanket over Earth's surface, hindering the loss of heat to space. The blanketing effect is greatest in the presence of low clouds and weakest for clear cold skies that contain little water vapour. Without this effect, the mean surface temperature of 15 °C (59 °F) would be some 30 °C colder. Conversely, as atmospheric concentrations of carbon dioxide, methane, chlorofluorocarbons (chlorofluorocarbon), and other absorbing gases continue to increase, in large part owing to human activities, surface temperatures should rise because of the capacity of such gases to trap infrared radiation. The exact amount of this temperature increase, however, remains uncertain because of unpredictable changes in other atmospheric components, especially cloud cover. An extreme example of such an effect (commonly dubbed the greenhouse effect) is that produced by the dense atmosphere of the planet Venus, which results in surface temperatures of about 475 °C (887 °F). This condition exists in spite of the fact that the high reflectivity of the Venusian clouds causes the planet to absorb less solar radiation than Earth.

Average radiation budgets
      The difference between the solar (heat) radiation absorbed and the thermal radiation (thermal energy) emitted to space determines Earth's radiation budget. Since there is no appreciable long-term trend in planetary temperature, it may be concluded that this budget is essentially zero on a global long-term average. Latitudinally, it has been found that much more solar radiation is absorbed (thermal conduction) at low latitudes than at high latitudes. On the other hand, thermal emission does not show nearly as strong a dependence on latitude, so the planetary radiation budget decreases systematically from the Equator to the poles. It changes from being positive to negative at latitudes of about 40° N and 40° S. The atmosphere and oceans, through their general circulation, act as vast heat engines, compensating for this imbalance by providing nonradiative mechanisms for the transfer of heat from the Equator to the poles.

      While Earth's surface absorbs a significant amount of thermal radiation because of the blanketing effect of the atmosphere, it loses even more through its own emission and thus experiences a net loss of long-wave radiation. This loss is only about 14 percent of the amount emitted by the surface and is less than the average gain of total absorbed solar energy. Consequently, the surface has on average a positive radiation budget.

      By contrast, the atmosphere emits thermal radiation both to space and to the surface, yet it receives long-wave radiation back from only the latter. This net loss of thermal energy cannot be compensated for by the modest gain of absorbed solar energy within the atmosphere. The atmosphere thus has a negative radiation budget, equal in magnitude to the positive radiation budget of the surface but opposite in sign. Nonradiative heat transfer again compensates for the imbalance, this time largely by vertical atmospheric motions involving the evaporation (vaporization) and condensation of water.

Surface-energy budgets
      The rate of temperature change in any region is directly proportional to the region's energy budget and inversely proportional to its heat capacity. While the radiation budget may dominate the average energy budget of many surfaces, nonradiative energy transfer and storage also are generally important when local changes are considered.

      Foremost among the cooling effects is the energy required to evaporate (vaporization) surface moisture, which produces atmospheric water vapour. Most of the latent heat contained in water vapour is subsequently released to the atmosphere during the formation of precipitating clouds, although a minor amount may be returned directly to the surface during dew or frost deposition. Evaporation increases with rising surface temperature, decreasing relative humidity, and increasing surface wind speed. transpiration by plants also increases evaporation rates, which explains why the temperature in an irrigated field is usually lower than that over a nearby dry road surface.

      Another important nonradiative mechanism is the exchange of heat that occurs when the temperature of the air is different from that of the surface. Depending on whether the surface is warmer or cooler than the air next to it, heat is transferred to or from the atmosphere by turbulent air motion (more loosely, by convection). This effect also increases with increasing temperature difference and with increasing surface wind speed. Direct heat transfer to the air may be an important cooling mechanism that limits the maximum temperature of hot dry surfaces. Alternatively, it may be an important warming mechanism that limits the minimum temperature of cold surfaces. Such warming is sensitive to wind speed, so calm conditions promote lower minimum temperatures.

      In a similar category, whenever a temperature difference occurs between the surface and the medium beneath the surface, there is a transfer of heat to or from the medium. In the case of land surfaces, heat is transferred by conduction (thermal conduction), a process where energy is conveyed through a material from one atom or molecule to another. In the case of water surfaces, the transfer is by convection and may consequently be affected by the horizontal transport of heat within large bodies of water.

 Average values of the different terms in the energy budgets of the atmosphere and surface are given in the diagram—>. The individual terms may be adjusted to suit local conditions and may be used as an aid to understanding the various temperature characteristics discussed in the next section.

Global variation of mean temperature
      Global variations of average surface-air temperatures are largely due to latitude, continentality, ocean currents (ocean current), and prevailing winds.

      The effect of latitude is evident in the large north-south gradients in average temperature that occur at middle and high latitudes in each winter hemisphere. These gradients are due mainly to the rapid decrease of available solar radiation but also in part to the higher surface reflectivity at high latitudes associated with snow and ice and low solar elevations. A broad area of the tropical ocean, by contrast, shows little temperature variation.

      Continentality is a measure of the difference between continental and marine climates and is mainly the result of the increased range of temperatures that occurs over land compared with water. This difference is a consequence of the much lower effective heat capacities of land surfaces as well as of their generally reduced evaporation rates. Heating or cooling of a land surface takes place in a thin layer, the depth of which is determined by the ability of the ground to conduct heat. The greatest temperature changes occur for dry, sandy soils, because they are poor conductors with very small effective heat capacities and contain no moisture for evaporation. By far the greatest effective heat capacities are those of water surfaces, owing to both the mixing of water near the surface and the penetration of solar radiation that distributes heating to depths of several metres. In addition, about 90 percent of the radiation budget of the ocean is used for evaporation. Ocean temperatures are thus slow to change.

      The effect of continentality may be moderated by proximity to the ocean, depending on the direction and strength of the prevailing winds. Contrast with ocean temperatures at the edges of each continent may be further modified by the presence of a north- or south-flowing ocean current. For most latitudes, however, continentality explains much of the variation in average temperature at a fixed latitude as well as variations in the difference between January and July temperatures.

Diurnal, seasonal, and extreme temperatures
      The diurnal range of temperature generally increases with distance from the sea and toward those places where solar radiation is strongest—in dry tropical climates and on high mountain plateaus (owing to the reduced thickness of the atmosphere to be traversed by the Sun's rays). The average difference between the day's highest and lowest temperatures is 3 °C (5 °F) in January and 5 °C (9 °F) in July in those parts of the British Isles nearest the Atlantic. The difference is 4.5 °C (8 °F) in January and 6.5 °C (12 °F) in July on the small island of Malta. At Tashkent, Uzbekistan, it is 9 °C (16 °F) in January and 15.5 °C (28 °F) in July, and at Khartoum, Sudan, the corresponding figures are 17 °C (31 °F) and 13.5 °C (24 °F). At Kandahār, Afghanistan, which lies more than 1,000 metres (about 3,300 feet) above sea level, it is 14 °C (25 °F) in January and 20 °C (36 °F) in July. There, the average difference between the day's highest and lowest temperatures exceeds 23 °C (41 °F) in September and October, when there is less cloudiness than in July. Near the ocean at Colombo, Sri L., the figures are 8 °C (14 °F) in January and 4.5 °C (8 °F) in July.

       World temperature extremesThe seasonal (season) variation of temperature and the magnitudes of the differences between the same month in different years and different epochs generally increase toward high latitudes and with distance from the ocean. Extreme temperatures observed in different parts of the world are listed in the table (World temperature extremes).

Variation with height
      There are two main levels where the atmosphere is heated—namely, at Earth's surface and at the top of the ozone layer (about 50 km, or 30 miles, up) in the stratosphere. Radiation balance shows a net gain at these levels in most cases. Prevailing temperatures tend to decrease with distance from these heating surfaces (apart from the ionosphere (ionosphere and magnetosphere) and the outer atmospheric layers, where other processes are at work). The world's average lapse rate of temperature (change with altitude) in the lower atmosphere is 0.6 to 0.7 °C per 100 metres (about 1.1 to 1.3 °F per 300 feet). Lower temperatures prevail with increasing height above sea level for two reasons: (1) because there is a less favourable radiation balance in the free air, and (2) because rising air—whether lifted by convection currents above a relatively warm surface or forced up over mountains—undergoes a reduction of temperature associated with its expansion as the pressure of the overlying atmosphere declines. This is the adiabatic lapse rate of temperature, which equals about 1 °C per 100 metres (about 2 °F per 300 feet) for dry air and 0.5 °C per 100 metres (about 1 °F per 300 feet) for saturated air, in which condensation (with liberation of latent heat) is produced by adiabatic cooling. The difference between these rates of change of temperature (and therefore density) of rising air currents and the state of the surrounding air determines whether the upward currents are accelerated or retarded—i.e., whether the air is unstable, so vertical convection with its characteristically attendant tall cumulus cloud and shower development is encouraged or whether it is stable and convection is damped down.

      For these reasons, the air temperatures observed on hills and mountains are generally lower than on low ground, except in the case of extensive plateaus, which present a raised heating surface (and on still, sunny days, when even a mountain peak is able to warm appreciably the air that remains in contact with it).

Circulation, currents, and ocean-atmosphere interaction
      The circulation of the ocean is a key factor in air temperature distribution. Ocean currents (ocean current) that have a northward or southward component, such as the warm Gulf Stream in the North Atlantic or the cold Peru (Humboldt) Current (Peru Current) off South America, effectively exchange heat between low and high latitudes. In tropical latitudes the ocean accounts for a third or more of the poleward heat transport; at latitude 50° N, the ocean's share is about one-seventh. In the particular sectors where the currents are located, their importance is of course much greater than these figures, which represent hemispheric averages.

      A good example of the effect of a warm current is that of the Gulf Stream in January, which causes a strong east-west gradient in temperatures across the eastern edge of the North American continent. The relative warmth of the Gulf Stream affects air temperatures all the way across the Atlantic, and prevailing westerlies extend the warming effect deep into northern Europe. As a result, January temperatures of Tromsø, Nor. (69°40′ N), for example, average 24 °C (43 °F) above the mean for that latitude. The Gulf Stream maintains a warming influence in July, but it is not as noticeable because of the effects of continentality.

      The ocean, particularly in areas where the surface is warm, also supplies moisture to the atmosphere. This in turn contributes to the heat budget of those areas in which the water vapour is condensed into clouds, liberating latent heat in the process. This set of events occurs frequently in high latitudes and in locations remote from the ocean where the moisture was initially taken up.

      The great ocean currents are themselves wind-driven—set in motion by the drag of the winds over vast areas of the sea surface, especially where the tops of waves increase the friction with the air above. At the limits of the warm currents, particularly where they abut directly upon a cold current—as at the left flank of the Gulf Stream in the neighbourhood of the Grand Banks off Newfoundland and at the subtropical and Antarctic convergences in the oceans of the Southern Hemisphere—the strong thermal gradients in the sea surface result in marked differences in the heating of the atmosphere on either side of the boundary. These temperature gradients tend to position and guide the strongest flow of the jet stream (see below Jet streams (climate)) in the atmosphere above and thereby influence the development and steering of weather systems.

 Interactions between the ocean and the atmosphere proceed in both directions. They also operate at different rates. Some interesting lag effects, which are of value in long-range weather forecasting, arise through the considerably slower circulation of the ocean. Thus, enhanced strength of the easterly trade winds (trade wind) over low latitudes of the Atlantic north and south of the Equator impels more water toward the Caribbean and the Gulf of Mexico, producing a stronger flow and greater warmth in the Gulf Stream approximately six months later. Anomalies in the position of the Gulf Stream– Labrador Current boundary, which produce a greater or lesser extent of warm water near the Grand Banks, so affect the energy supply to the atmosphere and the development and steering of weather systems from that region that they are associated with rather persistent anomalies of weather pattern over the British Isles and northern Europe. Anomalies in the equatorial Pacific and in the northern limit of the Kuroshio Current (Kuroshio) (also called the Japan Current) seem to have effects on a similar scale. Indeed, through their influence on the latitude of the jet stream and the wavelength (that is, the spacing of cold trough and warm ridge regions) in the upper westerlies, these ocean anomalies exercise an influence over the atmospheric circulation that spreads to all parts of the hemisphere.

      Sea-surface temperature anomalies that recur in the equatorial Pacific at variable intervals of two to seven years can sometimes produce major climatic perturbations. One such anomaly is known as El Niño (Spanish for “The Child”; it was so named by Peruvian fishermen who noticed its onset during the Christmas season).

      During an El Niño event, warm surface water flows eastward from the equatorial Pacific, in at least partial response to weakening of the equatorial easterly winds, and replaces the normally cold upwelling surface water off the coast of Peru and Ecuador that is associated with the northward propagation of the cold Peru Current. The change in sea-surface temperature transforms the coastal climate from arid to wet. The event also affects atmospheric circulation in both hemispheres and is associated with changes in precipitation in regions of North America, Africa, and the western Pacific. For further information, see the section El Niño/Southern Oscillation (ocean).

Short-term temperature changes
      Many interesting short-term temperature fluctuations also occur, usually in connection with local weather disturbances. The rapid passage of a mid-latitude cold front, for example, can drop temperatures by 10 °C (18 °F) in a few minutes and, if followed by the sustained movement of a cold air mass, by as much as 50 °C in 24 hours, with life-threatening implications for the unwary. Temperature increases of up to 40 °C in a few hours also are possible downwind of major mountain ranges when air that has been warmed by the release of latent heat on the windward side of a range is forced to descend rapidly on the other side (such a wind is variously called chinook, foehn, or Santa Ana). Changes of this kind, however, involve a wider range of meteorological processes than discussed in this section. For a more detailed treatment, see below Atmospheric pressure and wind (climate).

Hubert Horace Lamb Roger Davies

Atmospheric humidity and precipitation
      Atmospheric humidity, which is the amount of water vapour or moisture in the air, is another leading climatic element, as is precipitation. All forms of precipitation, including drizzle, rain, snow, ice crystals, and hail, are produced as a result of the condensation of atmospheric moisture that forms clouds (cloud) in which some of the particles, by growth and aggregation, attain sufficient size to fall from the clouds and reach the ground.

Atmospheric humidity
      At 30 °C (86 °F), 4 percent of the volume of the air may be occupied by water molecules, but, where the air is colder than −40 °C (−40 °F), less than one-fifth of 1 percent of the air molecules can be water. Although the water vapour content may vary from one air parcel to another, these limits can be set because vapour capacity is determined by temperature. temperature has profound effects upon some of the indexes of humidity, regardless of the presence or absence of vapour.

      The connection between an effect of humidity and an index of humidity requires simultaneous introduction of effects and indexes. Vapour in the air is a determinant of weather, because it first absorbs the thermal radiation that leaves and cools Earth's surface and then emits thermal radiation that warms the planet. Calculation of absorption and emission requires an index of the mass of water in a volume of air. Vapour also affects the weather because, as indicated above, it condenses into clouds and falls as rain or other forms of precipitation. Tracing the moisture-bearing air masses requires a humidity index that changes only when water is removed or added.

Humidity indexes

      Absolute humidity is the vapour concentration or density in the air. If mv is the mass of vapour in a volume of air, then absolute humidity dv is simply dv = mv/ V, in which V is the volume and dv is expressed in grams per cubic metre. This index indicates how much vapour a beam of radiation must pass through. The ultimate standard in humidity measurement is made by weighing the amount of water gained by an absorber when a known volume of air passes through it; this measures absolute humidity, which may vary from 0 gram per cubic metre in dry air to 30 grams per cubic metre (0.03 ounce per cubic foot) when the vapour is saturated at 30 °C. The dv of a parcel of air changes, however, with temperature or pressure even though no water is added or removed, because, as the gas equation states, the volume V increases with the absolute, or kelvin, temperature and decreases with the pressure.

      The meteorologist requires an index of humidity that does not change with pressure or temperature. A property of this sort will identify an air mass when it is cooled or when it rises to lower pressures aloft without losing or gaining water vapour. Because all the gases will expand equally, the ratios of the weight of water to the weight of dry air, or the dry air plus vapour, will be conserved during such changes and will continue identifying the air mass.

      The mixing ratio r is the dimensionless ratio r = mv/ ma, where ma is the mass of dry air, and the specific humidity q is another dimensionless ratio q = mv/ (ma + mv). Because mv is less than 3 percent of ma at normal pressure and temperatures cooler than 30 °C, r and q are practically equal. These indexes are usually expressed in grams per kilogram because they are so small; the values range from 0 grams per kilogram in dry air to 28 grams per kilogram in saturated air at 30 °C. Absolute and specific humidity indexes have specialized uses, so they are not familiar to most people.

       relative humidity (U) is so commonly used that a statement of humidity, without a qualifying adjective, can be assumed to be relative humidity. U can be defined, then, in terms of the mixing ratio r that was introduced above. U = 100r/ rw, which is a dimensionless percentage. The divisor rw is the saturation mixing ratio, or the vapour capacity. Relative humidity is therefore the water vapour content of the air relative to its content at saturation. Because the saturation mixing ratio is a function of pressure, and especially of temperature, the relative humidity is a combined index of the environment that reflects more than water content. In many climates the relative humidity rises to about 100 percent at dawn and falls to 50 percent by noon. A relative humidity of 50 percent may reflect many different quantities of vapour per volume of air or gram of air, and it will not likely be proportional to evaporation (vaporization).

      An understanding of relative humidity thus requires a knowledge of saturated vapour, which will be discussed later in the section on the relation between temperature and humidity. At this point, however, the relation between U and the absorption and retention of water from the air must be considered. Small pores retain water more strongly than large pores; thus, when a porous material is set out in the air, all pores larger than a certain size (which can be calculated from the relative humidity of the air) are dried out.

      The water content of a porous material at air temperature is fairly well indicated by the relative humidity. The complexity of actual pore sizes and the viscosity of the water passing through them makes the relation between U and moisture in the porous material imperfect and slowly achieved. The great suction also strains the walls of the capillaries, and the consequent shrinkage is used to measure relative humidity.

      The absorption of water by salt solutions is also related to relative humidity without much effect of temperature. The air above water saturated with sodium chloride is maintained at 75 to 76 percent relative humidity at a temperature between 0 and 40 °C (32 and 104 °F).

      In effect, relative humidity is a widely used environmental indicator, but U does respond drastically to changes in temperatures as well as moisture, a response caused by the effect of temperature upon the divisor rW in U.

Relation between temperature and humidity
      Tables that show the effect of temperature upon the saturation mixing ratio rw are readily available. Humidity of the air at saturation is expressed more commonly, however, as vapour pressure. Thus, it is necessary to understand vapour pressure and in particular the gaseous nature of water vapour.

      The pressure of the water vapour, which contributes to the pressure of the atmosphere, can be calculated from the absolute humidity dv by the gas equation:

      in which R is the gas constant, T the absolute temperature, Mw the molecular weight of water, and e the water vapour pressure in millibars (mb).

      Relative humidity can be defined as the ratio of the vapour pressure of a sample of air to the saturation pressure at the existing temperature. Further, the capacity for vapour and the effect of temperature can now be presented in the usual terms of saturation vapour pressure.

      Within a pool of liquid water, some molecules are continually escaping from the liquid into the space above, while more and more vapour molecules return to the liquid as the concentration of vapour rises. Finally, equal numbers are escaping and returning; the vapour is then saturated, and its pressure is known as the saturation vapour pressure, ew. If the liquid and vapour are warmed, relatively more molecules escape than return, and ew rises. There is also a saturation pressure with respect to ice. The vapour pressure curve of water has the same form as the curves for many other substances. Its location is fixed, however, by the boiling point of 100 °C (212 °F), where the saturation vapour pressure of water vapour is 1,013 mb (1 standard atmosphere), the standard pressure of the atmosphere at sea level. The decrease of the boiling point with altitude can be calculated. For example, the saturation vapour pressure at 40 °C (104 °F) is 74 mb (0.07 standard atmosphere), and the standard atmospheric pressure near 18,000 metres (59,000 feet) above sea level is also 74 mb; thus, it is where water boils at 40° C.

      The everyday response of relative humidity to temperature can be easily explained. On a summer morning, the temperature might be 15 °C (59 °F) and the relative humidity 100 percent. The vapour pressure would be 17 mb (0.02 standard atmosphere) and the mixing ratio about 11 parts per thousand (11 grams of water per kilogram of air by weight). During the day the air could warm to 25 °C (77 °F), while evaporation could add little water. At 25 °C the saturation pressure is fully 32 mb (0.03 standard atmosphere). If, however, little water has been added to the air, its vapour pressure will still be about 17 mb. Thus, with no change in vapour content, the relative humidity of the air has fallen from 100 to only 53 percent, illustrating why relative humidity does not identify air masses.

      The meaning of dew-point temperature can be illustrated by a sample of air with a vapour pressure of 17 mb. If an object at 15 °C is brought into the air, dew will form on the object. Hence, 15 °C is the dew-point temperature of the air—i.e., the temperature at which the vapour present in a sample of air would just cause saturation or the temperature whose saturation vapour pressure equals the present vapour pressure in a sample of air. Below freezing, this index is called the frost point. There is a one-to-one correspondence between vapour pressure and dew point. The dew point has the virtue of being easily interpreted because it is the temperature at which a blade of grass or a pane of glass will become wet with dew from the air. Ideally, it is also the temperature of fog or cloud formation.

      The clear meaning of dew point suggests a means of measuring humidity. A dew-point hygrometer (hygrometer) was invented in 1751. For this instrument, cold water was added to water in a vessel until dew formed on the vessel, and the temperature of the vessel, the dew point, provided a direct index of humidity. The greatest use of the condensation hygrometer has been to measure humidity in the upper atmosphere, where a vapour pressure of less than a thousandth millibar makes other means impractical.

      Another index of humidity, the saturation deficit, can also be understood by considering air with a vapour pressure of 17 mb. At 25 °C the air has (31 − 17), or 14, mb less vapour pressure than saturated vapour at the same temperature; that is, the saturation deficit is 14 mb (0.01 standard atmosphere).

      The saturation deficit has the particular utility of being proportional to the evaporation capability of the air. The saturation deficit can be expressed as

      and, because the saturation vapour pressure ew rises with rising temperature, the same relative humidity will correspond to a greater saturation deficit and evaporation at warm temperatures.

Humidity and climate
      The small amount of water in atmospheric vapour, relative to water on Earth, belies its importance. Compared with one unit of water in the air, the seas contain at least 100,000 units, the great glaciers 1,500, the porous earth nearly 200, and the rivers and lakes 4 or 5. The effectiveness of the vapour in the air is magnified, however, by its role in transferring water (hydrologic cycle) from sea to land by the media of clouds and precipitation and that in absorbing radiation.

      The vapour in the air is the invisible conductor that carries water from sea to land, making terrestrial life possible. Fresh water is distilled from the salt seas and carried over land by the wind. Water evaporates from vegetation, and rain falls on the sea too, but the sea is the bigger source, and rain that falls on land is most important to humans. The invisible vapour becomes visible near the surface as fog when the air cools to the dew point. The usual nocturnal cooling will produce fog patches in cool valleys. Or the vapour may move as a tropical air mass over cold land or sea, causing widespread and persistent fog, such as occurs over the Grand Banks off Newfoundland. The delivery of water by means of fog or dew is slight, however.

      When air is lifted, it is carried to a region of lower pressure, where it will expand and cool as described by the gas equation. It may rise up a mountain slope or over the front of a cooler, denser air mass. If condensation nuclei (condensation nucleus) are absent, the dew point may be exceeded by the cooling air, and the water vapour becomes supersaturated. If nuclei are present or if the temperature is very low, however, cloud droplets or ice crystals form, and the vapour is no longer in the invisible guise of atmospheric humidity.

      The invisible vapour has another climatic role—namely, absorbing and emitting radiation (solar radiation). The temperature of Earth and its daily variation are determined by the balance between incoming and outgoing radiation. The wavelength of the incoming radiation from the Sun is mostly shorter than 3 μm (0.0001 inch). It is scarcely absorbed by water vapour, and its receipt depends largely upon cloud cover. The radiation exchanged between the atmosphere and Earth's surface and the eventual loss to space is in the form of long waves. These long waves are strongly absorbed in the 3- to 8.5-μm band and in the greater than 11-μm range, where vapour is either partly or wholly opaque. As noted above, much of the radiation that is absorbed in the atmosphere is emitted back to Earth, and the surface receipt of long waves, primarily from water vapour and carbon dioxide in the atmosphere, is slightly more than twice the direct receipt of solar radiation at the surface. Thus, the invisible vapour in the atmosphere combines with clouds and the advection (horizontal movement) of air from different regions to control the surface temperature.

      The world distribution of humidity can be portrayed for different uses by different indexes. To appraise the quantity of water carried by the entire atmosphere, the moisture in an air column above a given point on Earth is expressed as a depth of liquid water. It varies from 0.5 mm (0.02 inch) over the Himalayas and 2 mm (0.08 inch) over the poles in winter to 8 mm (0.3 inch) over the Sahara, 54 mm (2 inches) in the Amazon region, and 64 mm (2.5 inches) over India during the wet season. During summer the air over the United States transports 16 mm (0.6 inch) of water vapour over the Great Basin and 45 mm (1.8 inches) over Florida.

      The humidity of the surface air may be mapped as vapour pressure, but a map of this variable looks much like that of temperature. Warm places are moist, and cool ones are dry; even in deserts the vapour pressure is normally 13 mb (0.01 standard atmosphere), whereas over the northern seas it is only about 4 mb (0.004 standard atmosphere). Certainly the moisture in materials in two such areas will be just the opposite, so relative humidity is a more widely useful index.

Average relative humidity
      The average relative humidity for July reveals the humidity provinces of the Northern Hemisphere when aridity is at a maximum. At other times the relative humidity generally will be higher. The humidities over the Southern Hemisphere in July indicate the humidities that comparable regions in the Northern Hemisphere will attain in January, just as July in the Northern Hemisphere suggests the humidities in the Southern Hemisphere during January. A contrast is provided by comparing a humid cool coast to a desert. The midday humidity on the Oregon coast, for example, falls only to 80 percent, whereas in the Nevada desert it falls to 20 percent. At night the contrast is less, with averages being over 90 and about 50 percent, respectively.

      Although the dramatic regular decrease of relative humidity from dawn to midday has been attributed largely to warming rather than declining vapour content, the content does vary regularly. In humid environments, daytime evaporation increases the water vapour content of the air, and the mixing ratio, which may be about 12 grams per kilogram, rises by 1 or 2 grams per kilogram in temperate places and may attain 16 grams per kilogram in a tropical rainforest. In arid environments, however, little evaporation moistens the air, and daytime turbulence tends to bring down dry air; this decreases the mixing ratio by as much as 2 grams per kilogram.

      Humidity also varies regularly with altitude. On the average, fully half the water in the atmosphere lies below 0.25 km (about 0.2 mile), and satellite observations over the United States in April revealed 1 mm (0.04 inch) or less of water in all the air above 6 km (4 miles). A cross section of the atmosphere along 75° W longitude shows a decrease in humidity with height and toward the poles. The mixing ratio is 16 grams per kilogram just north of the Equator, but it decreases to 1 gram per kilogram at 50° N latitude or 8 km (5 miles) above the Equator. The transparent air surrounding mountains in fair weather is very dry indeed.

      Closer to the ground, the water vapour content also changes with height in a regular pattern. When water vapour is condensing on Earth's surface at night, the content is greater aloft than at the ground; during the day the content is, in most cases, less aloft than at the ground because of evaporation.

Evaporation (vaporization) and humidity
 Evaporation, mostly from the sea and from vegetation, replenishes the humidity of the air. It is the change of liquid water into a gaseous state, but it may be analyzed as diffusion. The rate of diffusion, or evaporation, will be proportional to the difference between the pressure of the water vapour in the free air and the vapour that is next to, and saturated by, the evaporating liquid. If the liquid and air have the same temperature, evaporation is proportional to the saturation deficit. It is also proportional to the conductivity of the medium between the evaporator and the free air. If the evaporator is open water, the conductivity will increase with ventilation. But if the evaporator is a leaf, the diffusing water must pass through the still air within the minute pores between the water inside and the dry air outside. In this case the porosity may modify the conductivity more than ventilation.

      The temperature of the evaporator is rarely the same as the air temperature, however, because each gram of evaporation consumes about 600 calories (2,500 joules) and thus cools the evaporator. The availability of energy to heat the evaporator is therefore as important as the saturation deficit and conductivity of the air. Outdoors, some of this heat may be transferred from the surrounding air by convection, but much of it must be furnished by radiation. Evaporation is faster on sunny days than on cloudy ones not only because the sunny day may have drier air but also because the Sun warms the evaporator and thus raises the vapour pressure at the evaporator. In fact, according to the well-known Penman calculation of evaporation (an equation that considers potential evaporation as a function of humidity, wind speed, radiation, and temperature), this loss of water is essentially determined by the net radiation balance during the day.

Paul Edward Waggoner Ed.

  precipitation is one of the three main processes (evaporation (vaporization), condensation, and precipitation) that constitute the hydrologic cycle, the continual exchange of water between the atmosphere and Earth's surface. water evaporates from ocean, land, and freshwater surfaces, is carried aloft as vapour by the air currents, condenses to form clouds (cloud), and ultimately is returned to Earth's surface as precipitation. The average global stock of water vapour in the atmosphere is equivalent to a layer of water 2.5 cm (1 inch) deep covering the whole Earth. Because Earth's average annual rainfall is about 100 cm (39 inches), the average time that the water spends in the atmosphere, between its evaporation from the surface and its return as precipitation, is about 1/40 of a year, or about nine days. Of the water vapour that is carried at all heights across a given region by the winds, only a small percentage is converted into precipitation and reaches the ground in that area. In deep and extensive cloud systems, the conversion is more efficient, but even in thunderclouds the quantities of rain and hail released amount to only some 10 percent of the total moisture entering the storm.

      In the measurement of precipitation, it is necessary to distinguish between the amount—defined as the depth of precipitation (calculated as though it were all rain) that has fallen at a given point during a specified interval of time—and the rate or intensity, which specifies the depth of water that has fallen at a point during a particular interval of time. Persistent moderate rain, for example, might fall at an average rate of 5 mm per hour (0.2 inch per hour) and thus produce 120 mm (4.7 inches) of rain in 24 hours. A thunderstorm might produce this total quantity of rain in 20 minutes, but at its peak intensity the rate of rainfall might become much greater—perhaps 120 mm per hour (4.7 inches per hour), or 2mm (0.08 inch) per minute—for a minute or two.

      The amount of precipitation falling during a fixed period is measured regularly at many thousands of places on Earth's surface by rather simple rain gauges. Measurement of precipitation intensity requires a recording rain gauge, in which water falling into a collector of known surface area is continuously recorded on a moving chart or a magnetic tape. Investigations are being carried out on the feasibility of obtaining continuous measurements of rainfall over large catchment areas by means of radar.

 Apart from the trifling contributions made by dew, frost, and rime, as well as desalination plants, the sole source of fresh water for sustaining rivers, lakes, and all life on Earth is provided by precipitation from clouds (cloud). Precipitation is therefore indispensable and overwhelmingly beneficial to humankind, but extremely heavy rainfall can cause great harm: soil erosion, landslides (landslide), and flooding (flood). Hailstorm damage to crops, buildings, and livestock can prove very costly.

Origin of precipitation in clouds (cloud)

Cloud formation
 Clouds are formed by the lifting of damp air, which cools by expansion as it encounters the lower pressures existing at higher levels in the atmosphere. The relative humidity increases until the air has become saturated with water vapour, and then condensation occurs on any of the aerosol particles suspended in the air. A wide variety of these exist in concentrations ranging from only a few per cubic centimetre in clean maritime air to perhaps 1 million per cubic cm (16 million per cubic inch) in the highly polluted air of an industrial city. For continuous condensation leading to the formation of cloud droplets, the air must be slightly supersaturated. Among the highly efficient condensation nuclei (condensation nucleus) are sea-salt particles and the particles produced by combustion (e.g., natural forest fires and man-made fires). Many of the larger condensation nuclei over land consist of ammonium sulfate. These are produced by cloud and fog droplets absorbing sulfur dioxide (sulfur oxide) and ammonia from the air. Condensation onto the nuclei continues as rapidly as water vapour is made available through cooling; droplets about 10 μm (0.0004 inch) in diameter are produced in this manner. These droplets constitute a nonprecipitating cloud.

Cloud types
 The meteorologist classifies clouds mainly by their appearance, according to an international system similar to one proposed in 1803. But because the dimensions, shape, structure, and texture of clouds are influenced by the kind of air movements that result in their formation and growth and by the properties of the cloud particles, much of what was originally a purely visual classification can now be justified on physical grounds.

      The first International Cloud Atlas was published in 1896. Developments in aviation during World War I stimulated interest in cloud formations and in their importance as an aid in short-range weather forecasting. This led to the publication of a more extensive atlas, the International Atlas of Clouds and States of Sky, in 1932 and to a revised edition in 1939. After World War II, the World Meteorological Organization published a new International Cloud Atlas (1956) in two volumes. It contains 224 plates, describing 10 main cloud genera (families) subdivided into 14 species based on cloud shape and structure. Nine general varieties, based on transparency and geometric arrangement, also are described. The genera, listed according to their height, are as follows:

 1. High: mean heights from 5 to 13 km, or 3 to 8 miles (see photograph—>)

      a. Cirrus

      b. Cirrocumulus

      c. Cirrostratus

 2. Middle: mean heights 2 to 7 km, or 1 to 4 miles (see photograph—>)

      a. Altocumulus

      b. Altostratus

      c. Nimbostratus

 3. Low: mean heights 0 to 2 km, or 0 to 1.2 miles (see photograph—>)

      a. Stratocumulus

      b. Stratus

      c. Cumulus

      d. Cumulonimbus

      Heights given are approximate averages for temperate latitudes. Clouds of each genus are generally lower in the polar regions and higher in the tropics.

      Four principal classes are recognized when clouds are classified according to the kind of air motions that produce them: (1) layer clouds formed by the widespread regular ascent of air, (2) layer clouds formed by widespread irregular stirring or turbulence, (3) cumuliform clouds formed by penetrative convection, and (4) orographic clouds formed by the ascent of air over hills and mountains.

      The widespread layer clouds associated with cyclonic depressions (cyclone) (see below Cyclones and anticyclones (climate)), near fronts (front) and other inclement-weather systems, are frequently composed of several layers that may extend up to 9 km (5.6 miles) or more, separated by clear zones that become filled in as rain or snow develops. These clouds are formed by the slow, prolonged ascent of a deep layer of air, in which a rise of only a few centimetres per second is maintained for several hours. In the neighbourhood of fronts, vertical velocities become more pronounced and may reach about 10 cm (4 inches) per second.

      Most of the high cirrus clouds visible from the ground lie on the fringes of cyclonic cloud systems, and, though due primarily to regular ascent, their pattern is often determined by local wave disturbances that finally trigger their formation after the air has been brought near its saturation point by the large-scale lifting.

      On a cloudless night, the ground cools by radiating heat into space without heating the air adjacent to the ground. If the air were quite still, only a very thin layer would be chilled by contact with the ground. More usually, however, the lower layers of the air are stirred by motion over the rough ground, so the cooling is distributed through a much greater depth. Consequently, when the air is damp or the cooling is great, a fog a few hundred metres deep may form, rather than a dew produced by condensation on the ground.

      In moderate or strong winds, the irregular stirring near the surface distributes the cooling upward, and the fog may lift from the surface to become a stratus cloud, which is not often more than 600 metres (about 2,000 feet) thick.

      Radiational cooling from the upper surfaces of fogs and stratus clouds promotes an irregular convection within the cloud layer and causes the surfaces to have a waved or humped appearance. When the cloud layer is shallow, billows and clear spaces may develop; it is then described as stratocumulus instead of stratus.

      Usually, cumuliform clouds appearing over land are formed by the rise of discrete masses of air from near the sunlight-warmed surface. These rising lumps of air, or thermals (updraft and downdraft), may vary in diameter from a few tens to hundreds of metres as they ascend and mix with the cooler, drier air above them. Above the level of the cloud base, the release of latent heat of condensation tends to increase the buoyancy of the rising masses, which tower upward and emerge at the top of the cloud with rounded upper surfaces.

      At any moment a large cloud may contain a number of active thermals and the residues of earlier ones. A new thermal rising into a residual cloud will be partially protected from having to mix with the cool, dry environment and therefore may rise farther than its predecessor. Once a thermal has emerged as a cloud turret at the summit or the flanks of the cloud, rapid evaporation of the droplets chills the cloud borders, destroys the buoyancy, and produces sinking. A cumulus thus has a characteristic pyramidal shape and, viewed from a distance, appears to have an unfolding motion, with fresh cloud masses continually emerging from the interior to form the summit and then sinking aside and evaporating.

      In settled weather, cumulus clouds are well scattered and small; horizontal and vertical dimensions are only a kilometre or two. In disturbed weather, they cover a large part of the sky, and individual clouds may tower as high as 10 km (6 miles) or more, often ceasing their growth only upon reaching the stable stratosphere. These clouds produce heavy showers, hail, and thunderstorms .

 At the level of the cloud base, the speed of the rising air masses is usually about 1 metre (3.3 feet) per second but may reach 5 metres (16 feet) per second, and similar values are measured inside smaller clouds. The upcurrents in thunderclouds, however, often exceed 5 metres per second and may reach 30 metres (98 feet) per second or more.

      The rather special orographic clouds are produced by the ascent of air over hills and mountains. The air stream is set into oscillation when it is forced over the hill, and the clouds form in the crests of the (almost) stationary waves. There may thus be a succession of such clouds stretching downwind of the mountain, which remain stationary relative to the ground in spite of strong winds that may be blowing through the clouds. The clouds have very smooth outlines and are called lenticular (lens-shaped) or “wave” clouds. Thin wave clouds may form at great heights (up to 10 km, even over hills only a few hundred metres high) and occasionally are observed in the stratosphere (at 20 to 30 km [12 to 19 miles]) over the mountains of Norway, Scotland, Iceland, and Alaska. These atmospheric wave clouds are known as nacreous or “mother-of-pearl” clouds because of their brilliant iridescent colours.

Mechanisms of precipitation release
      Growing clouds are sustained by upward air currents (updraft and downdraft), which may vary in strength from a few centimetres per second to several metres per second. Considerable growth of the cloud droplets (with falling speeds of only about 1 cm, or 0.4 inch, per second) is therefore necessary if they are to fall through the cloud, survive evaporation in the unsaturated air below, and reach the ground as drizzle or rain. The production of a few large particles from a large population of much smaller ones may be achieved in one of two ways. The first of these depends on the fact that cloud droplets are seldom of uniform size; droplets form on nuclei of various sizes and grow under slightly different conditions and for different lengths of time in different parts of the cloud. A droplet appreciably larger than average will fall faster than the smaller ones and so will collide and fuse (coalesce) with some of those that it overtakes. Calculations show that, in a deep cloud containing strong upward air currents and high concentrations of liquid water, such a droplet will have a sufficiently long journey among its smaller neighbours to grow to raindrop size. This coalescence mechanism is responsible for the showers that fall in tropical and subtropical regions from clouds whose tops do not reach altitudes where air temperatures are below 0 °C (32 °F) and therefore cannot contain ice crystals. radar evidence also suggests that showers in temperate latitudes may sometimes be initiated by the coalescence of waterdrops, although the clouds may later reach heights at which ice crystals may form in their upper parts.

      The second method of releasing precipitation can operate only if the cloud top reaches elevations at which air temperatures are below 0 °C and the droplets in the upper cloud regions become supercooled. At temperatures below −40 °C (−40 °F), the droplets freeze automatically or spontaneously. At higher temperatures, they can freeze only if they are infected with special minute particles called ice nuclei. The origin and nature of these nuclei are not known with certainty, but the most likely source is clay-silicate particles carried up from the ground by the wind. As the temperature falls below 0 °C, more and more ice nuclei (freezing nucleus) become active, and ice crystals appear in increasing numbers among the supercooled droplets. Such a mixture of supercooled droplets and ice crystals is unstable, however. The cloudy air is usually only slightly supersaturated with water vapour with respect to the droplets and is strongly oversaturated with respect to ice crystals; the latter thus grow more rapidly than the droplets. After several minutes, the growing crystals acquire falling speeds of tens of centimetres per second, and several of them may become joined to form a snowflake. In falling into the warmer regions of the cloud, this flake may melt and hit ground as a raindrop.

      The deep, extensive, multilayer cloud systems, from which precipitation of a widespread persistent character falls, are generally formed in cyclonic depressions (lows) and near fronts. Cloud systems of this type are associated with feeble upcurrents of only a few centimetres per second that last for at least several hours. Although the structure of these great rain-cloud systems is being explored by aircraft and radar, it is not yet well understood. That such systems rarely produce rain, as distinct from drizzle, unless their tops are colder than about −12 °C (10 °F) suggests that ice crystals are mainly responsible. This view is supported by the fact that the radar signals from these clouds usually take a characteristic form that has been clearly identified with the melting of snowflakes.

Showers, thunderstorms (thunderstorm), and hail
 Precipitation from shower clouds and thunderstorms, whether in the form of raindrops, pellets of soft hail, or true hailstones, is generally of great intensity and shorter duration than that from layer clouds and is usually composed of larger particles. The clouds are characterized by their large vertical depth, strong vertical air currents, and high concentrations of liquid water, all factors favouring the rapid growth of precipitation elements by the accretion of cloud droplets.

      In a cloud composed wholly of liquid water, raindrops may grow by coalescence. For example, a droplet being carried up from the cloud base grows as it ascends by sweeping up smaller droplets. When it becomes too heavy to be supported by the upcurrents, the droplet falls, continuing to grow by the same process on its downward journey. Finally, if the cloud is sufficiently deep, the droplet will emerge from its base as a raindrop.

      In a dense, vigorous cloud several kilometres deep, the drop may attain its limiting stable diameter (about 6 mm [0.2 inch]) before reaching the cloud base and thus will break up into several large fragments. Each of these may continue to grow and attain breakup size. The number of raindrops may increase so rapidly in this manner that after a few minutes the accumulated mass of water can no longer be supported by the upcurrents and falls as a heavy shower. These conditions occur more readily in tropical regions. In temperate regions where the freezing level (0 °C) is much lower in elevation, conditions are more favourable for the ice-crystal mechanism.

      The hailstones that fall from deep, vigorous clouds in warm weather consist of a core surrounded by several alternate layers of clear and opaque ice. When the growing particle traverses a region of relatively high air temperature or high concentration of liquid water, or both, the transfer of heat from the hailstone to the air cannot occur rapidly enough to allow all of the deposited water to freeze immediately. This results in the formation of a wet coating of slushy ice, which may later freeze to form a layer of compact, relatively transparent ice. If the hailstone then enters a region of lower temperature and lower water content, the impacting droplets may freeze individually to produce ice of relatively low density with air spaces between the droplets. The alternate layers are formed as the stone passes through regions in which the combination of air temperature, liquid-water content, and updraft speed allows alternately wet and dry growth.

      It is held by some authorities that lightning is closely associated with the appearance of precipitation, especially in the form of soft hail, and that the charge and the strong electric fields are produced by ice crystals or cloud droplets striking and bouncing off the undersurfaces of the hail pellets. For a detailed discussion of electrical effects in clouds, see below thunderstorms (thunderstorm).

Types of precipitation

      Liquid precipitation in the form of very small drops, with diameters between 0.2 and 0.5 mm (0.008 and 0.02 inch) and terminal velocities between 70 and 200 cm per second (28 and 79 inches per second), is defined as drizzle. It forms by the coalescence of even smaller droplets in low-layer clouds containing weak updrafts of only a few centimetres per second. High relative humidity below the cloud base is required to prevent the drops from evaporating before reaching the ground; drizzle is classified as slight, moderate, or thick. Slight drizzle produces negligible runoff from the roofs of buildings, and thick drizzle accumulates at a rate in excess of 1 mm per hour (0.04 inch per hour).

rain and freezing rain
 Liquid waterdrops with diameters greater than those of drizzle constitute rain. Raindrops rarely exceed 6 mm (0.2 inch) in diameter because they become unstable when larger than this and break up during their fall. The terminal velocities of raindrops at ground level range from 2 metres per second (7 feet per second) for the smallest to about 10 metres per second (30 feet per second) for the largest. The smaller raindrops are kept nearly spherical by surface-tension forces, but, as the diameter surpasses about 2 mm (0.08 inch), they become increasingly flattened by aerodynamic forces. When the diameter reaches 6 mm, the undersurface of the drop becomes concave because of the airstream, and the surface of the drop is sheared off to form a rapidly expanding “bubble” or “bag” attached to an annular ring containing the bulk of the water. Eventually the bag bursts into a spray of fine droplets, and the ring breaks up into a circlet of millimetre-sized drops.

      Rain of a given intensity is composed of a spectrum of drop sizes, the average and median drop diameters being larger in rains of greater intensity. The largest drops, which have a diameter greater than 5 mm (0.2 inch), appear only in the heavy rains of intense storms.

      When raindrops fall through a cold layer of air (colder than 0 °C, or 32 °F) and become supercooled, freezing rain occurs. The drops may freeze on impact with the ground to form a very slippery and dangerous “glazed” ice that is difficult to see because it is almost transparent.

snow and sleet
  Snow in the atmosphere can be subdivided into ice crystals and snowflakes. Ice crystals generally form on ice nuclei (nucleation) at temperatures appreciably below the freezing point. Below −40 °C (−40 °F) water vapour can solidify without the presence of a nucleus. Snowflakes are aggregates of ice crystals that appear in an infinite variety of shapes, mainly at temperatures near the freezing point of water.

      In British terminology, sleet is the term used to describe precipitation of snow and rain together or of snow melting as it falls. In the United States, it is used to denote partly frozen ice pellets.

  Ice crystal shape and temperature at formationSnow crystals (crystal) generally have a hexagonal pattern, often with beautifully intricate shapes. Three- and 12-branched forms occur occasionally. The hexagonal form of the atmospheric ice crystals, their varying size and shape notwithstanding, is an outward manifestation of an internal arrangement in which the oxygen atoms form an open lattice (network) with hexagonally symmetrical structure. According to a recent internationally accepted classification, there are seven types of snow crystals: plates, stellars, columns, needles, spatial dendrites, capped columns, and irregular crystals. The size and shape of the snow crystals depend mainly on the temperature of their formation and on the amount of water vapour that is available for deposition. The two principal influences are not independent; the possible water vapour admixture of the air decreases strongly with decreasing temperature. The vapour pressure in equilibrium, or state of balance, with a level surface of pure ice is 50 times greater at −2 °C (28 °F) than at −42 °C (−44 °F), the likely limits of snow formation in the air. Crystal shape and temperature at formation are related in the table (Ice crystal shape and temperature at formation).

      At temperatures above about −40 °C (−40 °F), the crystals form on nuclei of very small size that float in the air (heterogeneous nucleation). The nuclei consist predominantly of silicate minerals (silicate mineral) of terrestrial origin, mainly clay minerals (clay mineral) and micas (mica). At still lower temperatures, ice may form directly from water vapour (homogeneous nucleation). The influence of the atmospheric water vapour depends mainly on its degree of supersaturation with respect to ice.

      If the air contains a large excess of water vapour, the snow particles will grow fast, and there may be a tendency for dendritic (branching) growth. With low temperature, the excess water vapour tends to be small, and the crystals remain small. In relatively dry layers, the snow particles generally have simple forms. Complicated forms of crystals will cling together with others to form snowflakes that consist occasionally of up to 100 crystals; the diameter of such flakes may be as large as 2.5 cm (1 inch). This process will be furthered if the crystals are near the freezing point and wet, possibly by collision with undercooled water droplets. If a crystal falls into a cloud with great numbers of such drops, it will sweep some of them up. Coming into contact with ice, they freeze and form an ice cover around the crystal. Such particles are called soft hail or graupel .

      Snow particles constitute the clouds of cirrus type—namely cirrus, cirrostratus, and cirrocumulus—and many clouds of alto type. Ice and snow clouds originate normally only at temperatures some degrees below the freezing point; they predominate at −20 °C (−4 °F). In temperate and low latitudes these clouds occur in the higher layers of the troposphere. In tropical regions they hardly ever occur below 4,570 metres (15,000 feet). On high mountains and particularly in polar regions, they can occur near the surface and may appear as ice fogs. If cold air near the ground is overlain by warmer air (a very common occurrence in polar regions, especially in winter), mixture at the border leads to supersaturation in the cold air. Small ice columns and needles, “diamond dust,” will be formed and will float down, glittering, even from a cloudless sky. In the coldest parts of Antarctica, where temperatures near the surface are below −50 °C (−58 °F) on the average and rarely above −30 °C (−22 °F), the formation of diamond dust is a common occurrence. The floating and falling ice crystals produce in the light of the Sun and the Moon the manifold phenomena of atmospheric optics, halos (halo), arcs, circles, mock suns, some coronas, and iridescent clouds. Most of the different optical appearances can be explained by the shapes of the crystals and their position with respect to the light source.

      Most of the moderate to heavy rain in temperate latitudes depends on the presence of ice and snow particles in clouds. In the free atmosphere, droplets of fluid water can be undercooled considerably; typical ice clouds originate mainly at a temperature near −20 °C. At an identical temperature below the freezing point, the water molecules are kept more firmly in the solid than in the fluid state. The equilibrium pressure of the gaseous phase is smaller in contact with ice than with water. At −20 °C, which is the temperature of the formation of typical ice clouds (cirrus), the equilibrium pressure with respect to undercooled water (relative humidity 100 percent) is 22 percent greater than the equilibrium pressure of the water vapour in contact with ice. Hence, with an excess of water vapour beyond the equilibrium state, the ice particles tend to incorporate more water vapour and to grow more rapidly than the water droplets.

      Being larger and so less retarded by friction, the ice particles fall more rapidly. In their fall they sweep up some water droplets, which on contact become frozen. Thus, a cloud layer originally consisting mainly of undercooled water with few ice crystals is transformed into an ice cloud. The development of the anvil shape at the top of a towering cumulonimbus cloud shows this transformation very clearly. The larger ice particles overcome more readily the rising tendency of the air in the cloud. Falling into lower levels they grow, aggregating with other crystals and possibly with waterdrops, melt, and form raindrops when near-surface temperatures permit.

 Solid precipitation in the form of hard pellets of ice that fall from cumulonimbus clouds is called hail. It is convenient to distinguish between three types of hail particles.

      The first is soft hail, or snow pellets, which are white opaque rounded or conical pellets as large as 6 mm (0.2 inch) in diameter. They are composed of small cloud droplets frozen together, have a low density, and are readily crushed.

      The second is small hail (ice grains or pellets), which are transparent or translucent pellets of ice that are spherical, spheroidal, conical, or irregular in shape, with diameters of a few millimetres. They may consist of frozen raindrops, of largely melted and refrozen snowflakes, or of snow pellets encased in a thin layer of solid ice.

      True hailstones, the third type, are hard pellets of ice, larger than 5 mm (0.2 inch) in diameter, that may be spherical, spheroidal, conical, discoidal, or irregular in shape and often have a structure of concentric layers of alternately clear and opaque ice. A moderately severe storm may produce stones a few centimetres in diameter, whereas a very severe storm may release stones with a maximum diameter of 10 cm (4 inches) or more. Large damaging hail falls most frequently in the continental areas of middle latitudes (e.g., in the Nebraska-Wyoming-Colorado area of the United States, in South Africa, and in northern India) but is rare in equatorial regions. Terminal velocities of hailstones range from about 5 metres (16 feet) per second for the smallest stones to perhaps 40 metres (130 feet) per second for stones 5 cm (2 inches) in diameter.

World distribution of precipitation

Regional and latitudinal distribution
 The yearly precipitation averaged over the whole Earth is about 100 cm (39 inches), but this is distributed very unevenly. The regions of highest rainfall are found in the equatorial zone and the monsoon area of Southeast Asia. Middle latitudes receive moderate amounts of precipitation, but little falls in the desert regions of the subtropics and around the poles.

 If Earth's surface were perfectly uniform, the long-term average rainfall would be distributed in distinct latitudinal bands, but the situation is complicated by the pattern of the global winds (wind), the distribution of land and sea, and the presence of mountains. Because rainfall results from the ascent and cooling of moist air, the areas of heavy rain indicate regions of rising air, whereas the deserts occur in regions in which the air is warmed and dried during descent. In the subtropics, the trade winds bring plentiful rain to the east coasts of the continents, but the west coasts tend to be dry. On the other hand, in high latitudes the west coasts are generally wetter than the east coasts. Rain tends to be abundant on the windward slopes of mountain ranges but sparse on the lee sides.

 In the equatorial belt, the trade winds from both hemispheres converge (intertropical convergence zone) and give rise to a general upward motion of air, which becomes intensified locally in tropical storms that produce very heavy rains in the Caribbean, the Indian and southwest Pacific oceans, and the China Sea and in thunderstorms that are especially frequent and active over the land areas. During the annual cycle, the doldrums move toward the summer hemisphere, so outside a central region near the Equator, which has abundant rain at all seasons, there is a zone that receives much rain in summer but a good deal less in winter.

      The dry areas of the subtropics—such as the desert regions of North Africa, the Arabian Peninsula, South Africa, Australia, and central South America—are due to the presence of semipermanent subtropical anticyclones (anticyclone) in which the air subsides and becomes warm and dry. These high-pressure belts tend to migrate with the seasons and cause summer dryness on the poleward side and winter dryness on the equatorward side of their mean positions (see below Cyclones and anticyclones (climate)). The easterly trade winds, having made a long passage over the warm oceans, bring plentiful rains to the east coasts of the subtropical landmasses, but the west coasts and the interiors of the continents, which are often sheltered by mountain ranges, are very dry.

      In middle latitudes, weather and rainfall are dominated by traveling depressions and fronts that yield a good deal of rain in all seasons and in most places except the far interiors of the Asian and North American continents. Generally, rainfall is more abundant in summer, except on the western coasts of North America, Europe, and North Africa, where it is higher during the winter.

      At high latitudes and especially in the polar regions, the low precipitation is caused partly by subsidence of air in the high-pressure belts and partly by the low temperatures. Snow or rain occur at times, but evaporation from the cold sea and land surfaces is slow, and the cold air has little capacity for moisture.

      The influence of oceans and continents on rainfall is particularly striking in the case of the Indian monsoon. During the Northern Hemisphere winter, cool dry air from the interior of the continent flows southward and produces little rain over the land areas. After the air has traveled some distance over the warm tropical ocean, however, it releases heavy shower rains over the East Indies. During the northern summer, when the monsoon blows from the southwest, rainfall is heavy over India and Southeast Asia. These rains are intensified where the air is forced to ascend over the windward slopes of the Western Ghats and the Himalayas.

      The combined effects of land, sea, mountains, and prevailing winds show up in South America. There the desert in southern Argentina is sheltered by the Andes from the westerly winds blowing from the Pacific Ocean, and the west-coast desert not only is situated under the South Pacific subtropical anticyclone but is also protected by the Andes against rain-bearing winds from the Atlantic.

Amounts and variability
      The long-term average amounts of precipitation for a season or a year give little information on the regularity with which rain may be expected, particularly for regions where the average amounts are small. For example, at Iquique, a city in northern Chile, four years once passed without rain, whereas the fifth year gave 15 mm (0.6 inch); the five-year average was therefore 3 mm (0.1 inch). Clearly, such averages are of little practical value, and the frequency distribution or the variability of precipitation also must be known.

      The variability of the annual rainfall is closely related to the average amounts. For example, over the British Isles, which have a very dependable rainfall, the annual amount varies by less than 10 percent above the long-term average value. A variability of less than 15 percent is typical of the mid-latitude cyclonic belts of the Pacific and Atlantic oceans and of much of the wet equatorial regions. In the interiors of the desert areas of Africa, Arabia, and Central Asia, however, the rainfall in a particular year may deviate from the normal long-term average by more than 40 percent. The variability for individual seasons or months may differ considerably from that for the year as a whole, but again the variability tends to be higher where the average amounts are low.

      The heaviest annual rainfall in the world was recorded at the village of Cherrapunji, India, where 26,470 mm (1,042 inches) fell between August 1860 and July 1861. The heaviest rainfall in a period of 24 hours was 1,870 mm (74 inches), recorded at the village of Cilaos, Réunion, in the Indian Ocean on March 15–16, 1952. The lowest recorded rainfall in the world occurred at Arica, a port city in northern Chile. An annual average, taken over a 43-year period, was only 0.5 mm (0.02 inch).

      Although past records give some guide, it is not possible to estimate very precisely the maximum possible precipitation that may fall in a given locality during a specified interval of time. Much will depend on a favourable combination of several factors, including the properties of the storm and the effects of local topography. Thus, it is possible only to make estimates that are based on analyses of past storms or on theoretical calculations that attempt to maximize statistically the various factors or the most effective combination of factors that are known to control the duration and intensity of the precipitation. For many important planning and design problems, however, estimates of the greatest precipitation to be expected at a given location within a specified number of years are required.

      In the designing of a dam, the highest 24-hour rainfall to be expected once in 30 years over the whole catchment area might be relevant. For dealing with such problems, a great deal of work has been devoted to determining from past records the frequency with which rainfalls of given intensity and total amount may be expected to reoccur at particular locations and also to determining the statistics of rainfall for a specific area from measurements made at only a few points.

Effects of precipitation

Raindrop impact and soil erosion
 Large raindrops, up to 6 mm (0.2 inch) in diameter, have terminal velocities of about 10 metres (30 feet) per second and so may cause considerable compaction and erosion of the soil by their force of impact. The formation of a compacted crust makes it more difficult for air and water to reach the roots of plants and encourages the water to run off the surface and carry away the topsoil with it. In hilly and mountainous areas, heavy rain may turn the soil into mud and slurry, which may produce enormous erosion by mudflow generation. Rainwater running off hard impervious surfaces or waterlogged soil may cause local flooding.

      The rainwater that is not evaporated or stored in the soil eventually runs off the surface and finds its way into rivers, streams, and lakes or percolates through the rocks and becomes stored in natural underground reservoirs. A given catchment area must achieve an overall balance such that precipitation (P) less evaporation of moisture from the surface (E) will equal storage in the ground (S) and runoff (R). This may be expressed: PE = S + R. The runoff may be determined by measuring the flow of water in the rivers with stream gauges, and the precipitation may be measured by a network of rain gauges, but storage and evaporation are more difficult to estimate.

      Of all the water that falls on Earth's surface, the relative amounts that run off, evaporate, or seep into the ground vary so much for different areas that no firm figures can be given for Earth as a whole. It has been estimated, however, that in the United States 10 to 50 percent of the rainfall runs off at once, 10 to 30 percent evaporates, and 40 to 60 percent is absorbed by the soil. Of the entire rainfall, 15 to 30 percent is estimated to be used by plants, either to form plant tissue or in transpiration.

Sir Basil John Mason Fritz P. Loewe Phillip J. Smith Ed.

Atmospheric pressure
       atmospheric pressure and wind are both significant controlling factors of Earth's (Earth) weather and climate. Although these two physical variables may at first glance appear to be quite different, they are in fact closely related. Wind exists because of horizontal and vertical differences (gradients) in pressure, yielding a correspondence that often makes it possible to use the pressure distribution as an alternative representation of atmospheric motions. Pressure is the force exerted on a unit area, and atmospheric pressure is equivalent to the weight of air above a given area on Earth's surface or within its atmosphere. This pressure is usually expressed in millibars (mb; 1 mb equals 1,000 dynes per square cm) or in kilopascals (kPa; 1 kPa equals 10,000 dynes per square cm). Distributions of pressure on a map are depicted by a series of curved lines called isobars (isobar), each of which connects points of equal pressure.

  At sea level the mean pressure is about 1,000 mb (100 kPa), varying by less than 5 percent from this value at any given location or time. Mean sea-level pressure values for the mid-winter months in the Northern Hemisphere are summarized in this first diagram—>, and mean sea-level pressure values for the mid-summer months are illustrated in the next diagram—>. Since charts of atmospheric pressure often represent average values over several days, pressure features that are relatively consistent day after day emerge, while more transient, short-lived features are removed. Those that remain are known as semipermanent pressure centres and are the source regions for major, relatively uniform bodies of air known as air masses (air mass). Warm, moist maritime tropical (mT) air forms over tropical and subtropical ocean waters in association with the high-pressure (anticyclone) regions prominent there. Cool, moist maritime polar (mP) air, on the other hand, forms over the colder subpolar ocean waters just south and east of the large, winter oceanic low-pressure (cyclone) regions. Over the continents, cold dry continental polar (cP) (polar air mass) air and extremely cold dry continental arctic (cA) air forms in the high-pressure regions that are especially pronounced in winter, while hot dry continental tropical (cT) air forms over hot desertlike continental domains in summer in association with low-pressure areas, which are sometimes called heat lows.

      A closer examination of the diagrams above reveals some interesting features. First, it is clear that sea-level pressure is dominated by closed high- and low-pressure centres, which are largely caused by differential surface heating between low and high latitudes and between continental and oceanic regions. High pressure tends to be amplified over the colder surface features. Second, because of seasonal changes in surface heating, the pressure centres exhibit seasonal changes in their characteristics. For example, the Siberian High (Siberian anticyclone), Aleutian Low, and Icelandic low that are so prominent in the winter virtually disappear in summer as the continental regions warm relative to surrounding bodies of water. At the same time, the Pacific and Atlantic highs amplify and migrate northward.

      At altitudes well above Earth's surface, the monthly average pressure distributions show much less tendency to form in closed centres but rather appear as quasi-concentric circles around the poles. This more symmetrical appearance reflects the dominant role of meridional (north-south) differences in radiative heating and cooling. Excess heating in tropical latitudes, in contrast to polar areas, produces higher pressure at upper levels in the tropics as thunderstorms (thunderstorm) transfer air to higher levels. In addition, the greater heating/cooling contrast in winter yields stronger pressure differences during this season. Perfect symmetry between the tropics and the poles is interrupted by wavelike atmospheric disturbances associated with migratory and semipermanent high- and low-pressure surface weather systems. These weather systems are most pronounced over the Northern Hemisphere, with its more prominent land-ocean contrasts and orographic (high-elevation) features.

Relationship of wind to pressure and governing forces
      The changing wind patterns are governed by Newton's (Newton, Sir Isaac) second law of motion, which states that the sum of the forces acting on a body equals the product of the mass of that body and the acceleration caused by those forces. The basic relationship between atmospheric pressure and horizontal wind is revealed by disregarding friction and any changes in wind direction and speed to yield the mathematical relationship

where u is the zonal wind speed (+ eastward), v the meridional wind speed (+ northward), f = 2ω sin ϕ (Coriolis parameter), ω the angular velocity of Earth's rotation, ϕ the latitude, ρ the air density (mass per unit volume), p the pressure, and x and y the distances toward the east and north, respectively. This simple non-accelerating flow is known as geostrophic balance and yields a motion field known as the geostrophic wind (geostrophic motion). Equation (1) expresses, for both the x and y directions, a balance between the force created by horizontal differences in pressure (the horizontal pressure-gradient force) and an apparent force that results from Earth's rotation (the Coriolis force). The pressure-gradient force expresses the tendency of pressure differences to effectuate air movement from higher to lower pressure. The Coriolis force arises because the air motions are observed on a rotating nearly spherical body. The total motion of a parcel of air has two parts: (1) the motion relative to Earth as if the planet were fixed, and (2) the motion given to the parcel of air by the planet's rotation. When the atmosphere is viewed from a fixed point in space, Earth's rotation is apparent. An observer in space would witness the total motion of the atmosphere. Conversely, an observer on the ground sees and measures only the relative motion of the atmosphere, because he is also rotating and cannot see directly the rotational motion applied by Earth. Instead, the observer on the ground sees the effect of the rotation as a deviation applied to the relative motion. The quantity that describes this deviation is the Coriolis force. Because the Coriolis force results from a ground-level frame of reference on a rotating planet, it is not a true force.

      More specifically, the observer on the ground experiences the Coriolis force as a deflection of the relative motion to the right in the Northern Hemisphere and to the left in the Southern Hemisphere. Of particular significance in this simple model of wind-pressure relationships is the fact that the geostrophic wind blows in a direction parallel to the isobars, with the low pressure on the observer's left as he looks downwind in the Northern Hemisphere and on his right in the Southern Hemisphere.

      Wind speed increases as the distance between isobars decreases (or pressure gradient increases). Curvature (i.e., changes in wind direction) can be added to this model with relative ease in a flow representation known as the gradient wind. The basic wind-pressure relationships, however, remain qualitatively the same. Of greatest importance is the fact that large-scale, observed winds tend to behave much as the geostrophic- or gradient-flow models predict in most of the atmosphere. The most notable exceptions occur in low latitudes, where the Coriolis parameter becomes very small—equation (1) cannot be used to provide a reliable wind estimate—and in the lowest kilometre of the atmosphere, where friction becomes important. The friction induced by airflow over the underlying surface reduces the wind speed and alters the simple balance of forces such that the wind blows with a component toward lower pressure.

Cyclones (cyclone) and anticyclones (anticyclone)
      Cyclones (cyclone) and anticyclones (anticyclone) are regions of relatively low and high pressure, respectively. They occur over most of Earth's surface in a variety of sizes ranging from the very large semipermanent examples described above to smaller, highly mobile systems. The latter are the focus of discussion in this section.

      Common to both cyclones and anticyclones are the characteristic circulation patterns. The geostrophic-wind and gradient-wind models dictate that, in the Northern Hemisphere, flow around a cyclone—cyclonic circulation—is counterclockwise, and flow around an anticyclone—anticyclonic circulation—is clockwise. Circulation directions are reversed in the Southern Hemisphere (see above the diagrams (climate) of mean sea-level pressure). In the presence of friction, the superimposed component of motion toward lower pressure produces a “spiraling” effect toward the low-pressure centre and away from the high-pressure centre.

      The cyclones that form outside the equatorial belt, known as extratropical cyclones (extratropical cyclone), may be regarded as large eddies (eddy) in the broad air currents that flow in the general direction from west to east around the middle and higher latitudes of both hemispheres (see below (climate)). They are an essential part of the mechanism by which the excess heat received from the Sun in Earth's equatorial belt is conveyed toward higher latitudes. These higher latitudes radiate more heat to space than they receive from the Sun, and heat must reach them by winds from the lower latitudes if their temperature is to be continually cool rather than cold. If there were no cyclones and anticyclones, the north-south movements of the air would be much more limited, and there would be little opportunity for heat to be carried poleward by winds of subtropical origin. Under such circumstances the temperature of the lower latitudes would increase, and the polar regions would cool; the temperature gradient between them would intensify.

      Strong horizontal gradients of temperature are particularly favourable for the formation and development of cyclones. The temperature difference between polar regions and the Equator builds up until it becomes sufficiently intense to generate new cyclones. As their associated cold fronts sweep equatorward and their warm fronts move poleward, the new cyclones reduce the temperature difference. Thus, the wind circulation on Earth represents a balance between the heating effects of solar radiation occurring in the polar regions and at the Equator. Wind circulation, through the effect of cyclones, anticyclones, and other wind systems, also periodically destroys this temperature contrast.

 Cyclones of a somewhat different character occur closer to the Equator, generally forming in latitudes between 10° to 30° N and S over the oceans. They generally are known as tropical cyclones (tropical cyclone) when their winds equal or exceed 74 miles (119 km) per hour. They are also known as hurricanes if they occur in the Atlantic Ocean and the Caribbean Sea, as typhoons in the western Pacific Ocean and the China Sea, and as cyclones off the coasts of Australia. These storms are of smaller diameter than the extratropical cyclones, ranging from 100 to 500 km (60 to 300 miles) in diameter, and are accompanied by winds that sometimes reach extreme violence. These storms are more fully described in the article tropical cyclone.

Extratropical cyclones (extratropical cyclone)
      Of the two types of large-scale cyclones, extratropical cyclones are the most abundant and exert influence on the broadest scale; they affect the largest percentage of Earth's surface. Furthermore, this class of cyclones is the principal cause of day-to-day weather changes experienced in middle and high latitudes and thus is the focal point of much of modern weather forecasting. The seeds for many current ideas concerning extratropical cyclones were sown between 1912 and 1930 by a group of Scandinavian meteorologists working in Bergen, Nor. This so-called Bergen school, founded by Norwegian meteorologist and physicist Vilhelm Bjerknes (Bjerknes, Vilhelm), formulated a model for a cyclone that forms as a disturbance along a zone of strong temperature contrast known as a front, which in turn constitutes a boundary between two contrasting air masses. In this model the masses of polar and mid-latitude air around the globe are separated by the polar front (the transition region separating warmer tropical air from colder polar air). This region possesses a strong temperature gradient, and thus it is a reservoir of potential energy that can be readily tapped and converted into the kinetic energy associated with extratropical cyclones.

  For this reservoir to be tapped, a cyclone (called a wave, or frontal, cyclone) must develop much in the way shown in the diagram. The feature that is of primary importance prior to cyclone development (cyclogenesis) is a front, represented in the initial stage (A) as a heavy black line with alternating triangles or semicircles attached to it. This stationary or very slow-moving front forms a boundary between cold and warm air and thus is a zone of strong horizontal temperature gradient (sometimes referred to as a baroclinic zone). Cyclone development is initiated as a disturbance along the front, which distorts the front into the wavelike configuration (B; wave appearance). As the pressure within the disturbance continues to decrease, the disturbance assumes the appearance of a cyclone and forces poleward and equatorward movements of warm and cold air, respectively, which are represented by mobile frontal boundaries. As depicted in the cyclonic circulation stage (C), the front that signals the advancing cold air ( cold front) is indicated by the triangles, while the front corresponding to the advancing warm air (warm front) is indicated by the semicircles. As the cyclone continues to intensify, the cold dense air streams rapidly equatorward, yielding a cold front with a typical slope of 1 to 50 and a propagation speed that is often 8 to 15 metres per second (about 18 to 34 miles per hour) or more. At the same time, the warm less-dense air moving in a northerly direction flows up over the cold air east of the cyclone to produce a warm front with a typical slope of 1 to 200 and a typically much slower propagation speed of about 2.5 to 8 metres per second (6 to 18 miles per hour). This difference in propagation speeds between the two fronts allows the cold front to overtake the warm front and produce yet another, more complicated frontal structure, known as an occluded front. An occluded front (D) is represented by a line with alternating triangles and semicircles on the same side. This occlusion process may be followed by further storm intensification. The separation of the cyclone from the warm air toward the Equator, however, eventually leads to the storm's decay and dissipation (E) in a process called cyclolysis.

 The life cycle of such an event is typically several days, during which the cyclone may travel from several hundred to a few thousand kilometres. In its path and wake occur dramatic weather changes. A typical sequence of weather possibly resulting from the approach and passage of a cyclone and its fronts through an area is depicted in the diagram—>. Shown in the occluded-front stage of the cyclogenesis diagram is a cross section of the clouds (cloud) and precipitation that usually occur along line ab. Warm frontal weather is most frequently characterized by stratiform clouds, which ascend as the front approaches and potentially yield rain or snow. The passing of a warm front brings a rise in air temperature and clearing skies. The warmer air, however, may also harbour the ingredients for rain shower or thunderstorm formation, a condition that is enhanced as the cold front approaches.

 The passage of the cold front is marked by the influx of colder air, the formation of stratocumulus clouds with some lingering rain or snow showers, and then eventual clearing. While this is an oft-repeated scenario, it is important to recognize that many other weather sequences can also occur. For example, the stratiform clouds of a warm front may have imbedded cumulus formations and thunderstorms; the warm sector might be quite dry and yield few or no clouds; the pre-cold-front weather may closely resemble that found ahead of the warm front; or the post-cold-front air may be completely cloud-free. Cloud patterns oriented along fronts and spiraling around the cyclone vortex are consistently revealed in satellite pictures of Earth.

      The actual formation of any area of low pressure requires that mass in the column of air lying above Earth's surface be reduced. This loss of mass then reduces the surface pressure. In the late 1930s and early '40s, three members of the Bergen school—Norwegian American meteorologists Jacob Bjerknes (Bjerknes, Jacob) and Jørgen Holmboe and Swedish American meteorologist Carl-Gustaf Rossby (Rossby, Carl-Gustaf Arvid)—recognized that transient surface disturbances were accompanied by complementary wave features in the flow in the middle and higher atmospheric layers associated with the jet stream. These wave features are accompanied by regions of mass divergence and convergence that support the growth of surface-pressure fields and direct their movement.

 While extratropical cyclones form and intensify in association with fronts, there are small-scale cyclones that appear in the middle of a single air mass. A notable example is a class of cyclones, generally smaller than the frontal variety, that form in polar air streams in the wake of a frontal cyclone. These so-called polar lows are most prominent in subpolar marine environments and are thought to be caused by the transfer of heat and moisture from the warmer water surface into the overlying polar air and by supporting middle-tropospheric circulation features. Other small-scale cyclones form on the lee side of mountain barriers as the general westerly flow is disturbed by the mountain. These “lee cyclones” may produce major windstorms (windstorm) and dust storms downstream of a mountain barrier.

      While cyclones are typically regions of inclement weather, anticyclones are usually meteorologically quiet regions. Generally larger than cyclones, anticyclones exhibit persistent downward motions and yield dry stable air that may extend horizontally many hundreds of kilometres.

      In most cases, an actively developing anticyclone forms over a ground location in the region of cold air behind a cyclone as it moves away. This anticyclone forms before the next cyclone advances into the area. Such an anticyclone is known as a cold anticyclone. A result of the downward air motion in an anticyclone, however, is compression of the descending air. As a consequence of this compression, the air is warmed. Thus, after a few days, the air composing the anticyclone at levels 2 to 5 km (1 to 3 miles) above the ground tends to increase in temperature, and the anticyclone is transformed into a warm anticyclone.

      Warm anticyclones move slowly, and cyclones are diverted around their periphery. During their transformation from cold to warm status, anticyclones usually move out of the main belt followed by cyclones in middle latitudes and often amalgamate with the quasi-permanent bands of relatively high pressure found in both hemispheres around latitude 20° to 30°—the so-called subtropical anticyclones (subtropical high). On some occasions the warm anticyclones remain in the belt normally occupied by the mid-latitude westerly winds. The normal cyclone tracks are then considerably modified; atmospheric depressions (areas of low pressure) are either blocked in their eastward progress or diverted to the north or south of the anticyclone. Anticyclones that interrupt the normal circulation of the westerly wind in this way are called blocking anticyclones, or blocking highs. They frequently persist for a week or more, and the occurrence of a few such blocking anticyclones may dominate the character of a season. Blocking anticyclones are particularly common over Europe, the eastern Atlantic, and the Alaskan area.

      The descent and warming of the air in an anticyclone might be expected to lead to the dissolution of clouds and the absence of rain. Near the centre of the anticyclone, the winds are light and the air can become stagnant. Air pollution can build up as a result. The city of Los Angeles, for example, often has poor air quality because it is frequently under a stationary anticyclone. In winter the ground cools, and the lower layers of the atmosphere also become cold. fog may be formed as the air is cooled to its dew point in the stagnant air. Under other circumstances, the air trapped in the first kilometre above Earth's surface may pick up moisture from the sea or other moist surfaces, and layers of cloud may form in areas near the ground up to a height of about 1 km (0.6 mile). Such layers of cloud can be persistent in anticyclones (except over the continents in summer), but they rarely grow thick enough to produce rain. If precipitation occurs, it is usually drizzle or light snow.

      Anticyclones are often regions of clear skies and sunny weather in summer; at other times of the year, cloudy and foggy weather—especially over wet ground, snow cover, and the ocean—may be more typical. Winter anticyclones produce colder than average temperatures at the surface, particularly if the skies remain clear. Anticyclones are responsible for periods of little or no rain, and such periods may be prolonged in association with blocking highs.

Cyclone and anticyclone climatology
  Migrating cyclones and anticyclones tend to be distributed around certain preferred regions, known as tracks, that emanate from preferred cyclogenetic and anticyclogenetic regions. Favoured cyclogenetic regions in the Northern Hemisphere are found on the lee side of mountains and off the east coasts of continents. Cyclones then track east or southeast before eventually turning toward the northeast and decaying. The tracks are displaced farther northward in July, reflecting the more northward position of the polar front in summer. Continental cyclones usually intensify at a rate of 0.5 mb (0.05 kPa) per hour or less, although more dramatic examples can be found. Marine cyclones, on the other hand, often experience explosive development in excess of 1 mb (0.1 kPa) per hour, particularly in winter.

      Anticyclones tend to migrate equatorward out of the cold air mass regions and then eastward before decaying or merging with a warm anticyclone. Like cyclones, warm anticyclones also slowly migrate poleward with the warm season.

      In the Southern Hemisphere, where most of Earth's surface is covered by oceans, the cyclones are distributed fairly uniformly through the various longitudes. Typically, cyclones form initially in latitudes 30° to 40° S and move in a generally southeastward direction, reaching maturity in latitudes near 60° S. Thus, the Antarctic continent is usually ringed by a number of mature or decaying cyclones. The belt of ocean from 40° to 60° S is a region of persistent, strong westerly winds that form part of the circulation to the north of the main cyclone centres; These are the “ roaring forties,” where the westerly winds are interrupted only at intervals by the passage southeastward of developing cyclones.

Local winds

Scale classes
      Organized wind systems occur in spatial dimensions ranging from tens of metres to thousands of kilometres and possess residence times that vary from seconds to weeks. The concept of scale considers the typical size and lifetime of a phenomenon. Since the atmosphere exhibits such a large variety of both spatial and temporal scales, efforts have been made to group various phenomena into scale classes. The class describing the largest and longest-lived of these phenomena is known as the planetary scale. Such phenomena are typically a few thousand kilometres in size and have lifetimes ranging from several days to several weeks. Examples of planetary-scale phenomena include the semipermanent pressure centres discussed above and certain globe-encircling upper-air waves (see below Upper-air waves (climate)).

      A second class is known as the synoptic scale. Spanning smaller distances, a few hundred to a few thousand kilometres, and possessing shorter lifetimes, a few to several days, this class contains the migrating cyclones and anticyclones that control day-to-day weather changes. Sometimes the planetary and synoptic scales are combined into a single classification termed the large-scale, or macroscale. Large-scale wind systems are distinguished by the predominance of horizontal motions over vertical motions and by the preeminent importance of the Coriolis force in influencing wind characteristics. Examples of large-scale wind systems include the trade winds (trade wind) and the westerlies (Pacific Ocean).

      There is a third class of phenomena of even smaller size and shorter lifetime. In this class, vertical motions may be as significant as horizontal movement, and the Coriolis force often plays a less important role. Known as the mesoscale, this class is characterized by spatial dimensions of ten to a few hundred kilometres and lifetimes of a day or less. Because of the shorter time scale and because the other forces may be much larger, the effect of the Coriolis force in mesoscale phenomena is sometimes neglected.

      Two of the best-known examples of mesoscale phenomena are the thunderstorm and its devastating by-product, the tornado (see thunderstorm; tornado). The present discussion focuses on less intense, though nevertheless commonly observed, wind systems that are found in rather specific geographic locations and thus are often referred to as local wind systems.

Local wind systems
 The so-called sea and land breeze circulation is a local wind system typically encountered along coastlines adjacent to large bodies of water and is induced by differences that occur between the heating or cooling of the water surface and the adjacent land surface. Water has a higher heat capacity (i.e., more units of heat are required to produce a given temperature change in a volume of water) than do the materials in the land surface. Daytime solar radiation penetrates to several metres into the water, the water vertically mixes, and the volume is slowly heated. In contrast, daytime solar radiation heats the land surface more quickly because it does not penetrate more than a few centimetres below the land surface. The land surface, now at a higher temperature relative to the air adjacent to it, transfers more heat to its overlying air mass and creates an area of low pressure. Thus, a circulation cell much like that depicted in the diagram—> is induced.It should be noted that the surface flow is from the water toward the land and thus is called a sea breeze.

      Since the landmass possesses a lower heat capacity than water, the land cools more rapidly at night than does the water. Consequently, at night the cooler landmass yields a cooler overlying air mass and creates a zone of relatively higher pressure. This produces a circulation cell with air motions opposite to those found during the day. This flow from land to water is known as a land breeze. The land breeze is typically shallower than the sea breeze since the cooling of the atmosphere over land is confined to a shallower layer at night than the heating of the air during the day.

      Sea and land breezes occur along the coastal regions of oceans or large lakes in the absence of a strong large-scale wind system during periods of strong daytime heating or nighttime cooling. Those who live within 10 to 20 km (6 to 12 miles) of the coastline often experience the cooler 19- to 37-km-per-hour (12- to 23-mile-per-hour) winds of the sea breeze on a sunny afternoon only to find it turn into a sultry land breeze late at night. One of the features of the sea and land breeze is a region of low-level air convergence in the termination region of the surface flow. Such convergence often induces local upward motions and cloud formations. Thus, in sea and land breeze regions, it is not uncommon to see clouds lying off the coast at night; these clouds are then dissipated by the daytime sea breeze, which forms new clouds, perhaps with showers occurring over land in the afternoon.

 Another group of local winds is induced by the presence of mountain and valley features on Earth's surface. One subset of such winds, known as mountain winds or breezes, is induced by differential heating or cooling along mountain slopes. During the day, solar heating of the sunlit slopes causes the overlying air to move upslope. These winds are also called anabatic flow (anabatic wind). At night, as the slopes cool, the direction of airflow is reversed, and cool downslope drainage motion occurs. Such winds may be relatively gentle or may occur in strong gusts, depending on the topographic configuration. These winds are one type of katabatic flow (katabatic wind). In an enclosed valley, the cool air that drains into the valley may give rise to a thick fog condition. Fog persists until daytime heating reverses the circulation and creates clouds associated with the upslope motion at the mountain top.

      Another subset of katabatic flow, called foehn winds (also known as chinook winds (chinook) east of the Rocky Mountains and as Santa Ana winds in southern California), is induced by adiabatic (adiabatic process) temperature changes occurring as air flows over a mountain. Adiabatic temperature changes are those that occur without the addition or subtraction of heat; they occur in the atmosphere when bundles of air are moved vertically. When air is lifted, it enters a region of lower pressure and expands. This expansion is accompanied by a reduction of temperature (adiabatic cooling). When air subsides, it contracts and experiences adiabatic warming. As air ascends on the windward side of the mountain, its cooling rate may be moderated by heat that is released during the formation of precipitation. However, having lost much of its moisture, the descending air on the leeward side of the mountain adiabatically warms faster than it was cooled on the windward ascent. Thus, the effect of this wind, if it reaches the surface, is to produce warm, dry conditions. Usually, such winds are gentle and produce a slow warming. On occasion, however, foehn winds may exceed 185 kilometres (115 miles) per hour and produce air-temperature increases of tens of degrees (sometimes more than 20 °C [36 °F]) within only a few hours.

      Other types of katabatic wind can occur when the underlying geography is characterized by a cold plateau adjacent to a relatively warm region of lower elevation. Such conditions are satisfied in areas in which major ice sheets or cold elevated land surfaces border warmer large bodies of water. Air over the cold plateau cools and forms a large dome of cold dense air. Unless held back by background wind conditions, this cold air will spill over into the lower elevations with speeds that vary from gentle (a few kilometres per hour) to intense (93 to 185 km [58 to 115 miles] per hour), depending on the incline of the slope of the terrain and the distribution of the background pressure field. Two special varieties of katabatic wind are well known in Europe. One is the bora, which blows from the highlands of Croatia, Bosnia and Herzegovina, and Montenegro to the Adriatic Sea; the other is the mistral, which blows out of central and southern France to the Mediterranean Sea. Creating blizzard conditions, intense katabatic winds often blow northward off the Antarctic Ice Sheet.

Zonal surface winds
  The diagrams of January—> and July—> mean sea-level pressure reveal that, on the average, certain geographic locations can expect to experience winds that emanate from one prevailing direction largely dictated by the presence of major semipermanent pressure systems. Such prevailing winds have long been known in marine environments because of their influence on the great sailing ships.

 Tropical and subtropical regions are characterized by a general band of low pressure lying near the Equator. This band is bounded by centres of high pressure that may extend poleward into the middle latitudes. Between these low- and high-pressure regions is the region of the tropical winds. Of these the most extensive are the trade winds (trade wind). So named because of their favourable influence on trade ships traveling across the subtropical North Atlantic, trade winds flow westward and somewhat in the direction of the Equator on the equatorward side of the subtropical high-pressure centres. The “root of the trades,” occurring on the eastern side of a subtropical high-pressure centre (subtropical high), is characterized by subsiding air. This produces the very warm, dry conditions above a shallow layer of oceanic stratus clouds found in the eastern extremes of the subtropical Atlantic and Pacific ocean basins. As the trade winds progress westward, however, subsidence abates, the air mass becomes more humid, and scattered showers appear. These showers occur particularly on islands with elevated terrain features that interrupt the flow of the warm moist air. The equatorward flow of the trade winds of the Northern and Southern hemispheres often results in a convergence of the two air streams in a region known as the intertropical convergence zone (ITCZ) (intertropical convergence zone). Deep convective clouds, showers, and thunderstorms occur along the ITCZ.

      When the air reaches the western extreme of the high-pressure centre, it turns poleward and then eventually returns eastward in the middle latitudes. The poleward-moving air is now warm and laden with moist maritime tropical air (mT); it gives rise to the warm, humid, showery climate characteristic of the Caribbean region, eastern South America, and the western Pacific island chains. The westerlies are associated with the changeable weather common to the middle latitudes. Migrating extratropical cyclones and anticyclones associated with contrasting warm moist air moving poleward from the tropics and cold dry air moving equatorward from polar latitudes yield periods of rain (sometimes with violent thunderstorms) (thunderstorm), snow, sleet, or freezing rain interrupted by periods of dry, sunny, and sometimes bitterly cold conditions. Furthermore, these patterns are seasonally dependent, with more intense cyclones and colder air prevailing in winter but with a higher incidence of thunderstorms common in spring and summer. In addition, these migrations and the associated climate are complicated by the presence of landmasses and major mountain features, particularly in the Northern Hemisphere.

      The westerlies lie on the equatorward side of the semipermanent subpolar centres of low pressure. Poleward of these centres, the surface winds turn westward again over significant portions of the subpolar latitudes. As in the middle latitudes, the presence of major landmasses, notably in the Northern Hemisphere, results in significant variations in these polar easterlies. In addition, the wind systems and the associated climate are seasonally dependent. During the short summer season, the wind systems of the polar latitudes are greatly weakened. During the long winter months, these systems strengthen, and periods of snow alternate with long intervals of dry cold air characteristic of continental polar or continental arctic air masses.

 These major regions of surface circulation and their associated pressure fields are related to mean meridional (north-south) circulation patterns as well (see the diagram—>). Although their presence is discernible in long-term mean statistics accumulated over a hemisphere, such cells are often difficult to detect on a daily basis at any given longitude.

Phillip J. Smith

Monsoons (monsoon)
   As shown in the January—> and July—> maps of sea-level atmospheric pressure, particularly strong seasonal pressure variations occur over continents. Such seasonal fluctuations, commonly called monsoons (monsoon), are more pronounced over land surfaces because these surfaces are subject to more significant seasonal temperature variations than water bodies. Since land surfaces both warm and cool faster than water bodies, they often quickly modify the temperature and density characteristics of air parcels passing over them.

      Monsoons blow for approximately six months from the northeast and six months from the southwest, principally in South Asia and parts of Africa. Summer monsoons have a dominant westerly component and a strong tendency to converge, rise, and produce rain. Winter monsoons have a dominant easterly component and a strong tendency to diverge, subside, and cause drought. Both are the result of differences in annual temperature trends over land and sea.

The Indian monsoon

 At the Equator the area near India is unique in that dominant or frequent westerly winds (Pacific Ocean) occur at the surface almost constantly throughout the year; the surface easterlies reach only to latitudes near 20° N in February, and even then they have a very strong northerly component. They soon retreat northward, and drastic changes take place in the upper-air circulation (see Jet streams (climate)). This is a time of transition between the end of one monsoon and the beginning of the next. Late in March, the high-sun season reaches the Equator and moves farther north. With it go atmospheric instability, convectional (that is, rising and turbulent) clouds (cloud), and rain. The westerly subtropical jet stream still controls the flow of air across northern India, and the surface winds (wind) are northeasterlies.

      As the high-sun season moves northward during April, India becomes particularly prone to rapid heating because the highlands to the north protect it from any incursions of cold air. There are three distinct areas of relative upper tropospheric (troposphere) warmth—namely, (1) above the southern Bay of Bengal (Bengal, Bay of), (2) above the Plateau of Tibet (Tibet, Plateau of), and (3) across the trunks of the various peninsulas that are relatively dry during this time. These three areas combine to form a vast heat-source region. The relatively warm area above the southern Bay of Bengal occurs mostly at the 500–100- millibar level. (This atmospheric pressure region typically occurs at elevations between 18,000 and 53,000 feet but may vary according to changes in heating and cooling.) It does not appear at a lower level and is probably caused by the release of condensation heat (associated with the change from water vapour to liquid water) at the top of towering cumulonimbus clouds along the advancing intertropical convergence (intertropical convergence zone). In contrast, a heat sink appears over the southern Indian Ocean as the relatively cloud-free air cools by emitting long-wavelength radiation. Monsoon winds at the surface blow from heat sink to heat source. As a result, by May the southwest monsoon is well-established over Sri Lanka, an island off the southeastern tip of the Indian peninsula.

      Also in May, the dry surface of Tibet (above 4,000 metres [13,100 feet]) absorbs and radiates heat that is readily transmitted to the air immediately above. At about 6,000 metres (19,700 feet) an anticyclonic (anticyclone) cell arises, causing a strong easterly flow in the upper troposphere above northern India. The subtropical jet stream suddenly changes its course to the north of the anticyclonic ridge and the highlands, though it may occasionally reappear southward of them for very brief periods. This change of the upper tropospheric circulation above northern India from westerly jet to easterly flow coincides with a reversal of the vertical temperature and pressure gradients between 600 and 300 millibars. On many occasions the easterly wind aloft assumes jet force. It anticipates by a few days the “burst,” or onset, of the surface southwesterly monsoon some 1,500 km (900 miles) farther south, with a definite sequential relationship, although the exact cause is not known. Because of India's inverted triangular shape, the land is heated progressively as the Sun moves northward. This accelerated spread of heating, combined with the general direction of heat being transported by winds, results in a greater initial monsoonal activity over the Arabian Sea (at late springtime), where a real frontal situation often occurs, than over the Bay of Bengal. The relative humidity (humidity) of coastal districts in the Indian region rises above 70 percent, and some rain occurs. Above the heated land, the air below 1,500 metres (5,000 feet) becomes unstable, but it is held down by the overriding easterly flow. This does not prevent frequent thunderstorms (thunderstorm) from occurring in late May.

      During June the easterly jet becomes firmly established at 150 to 100 millibars, an atmospheric pressure region typically occurring at elevations between 45,000 and 53,000 feet. It reaches its greatest speed at its normal position to the south of the anticyclonic ridge, at about 15° N from China through India.

      In Arabia (Saudi Arabia) it decelerates and descends to the middle troposphere (3,000 metres [9,800 feet]). A stratospheric (stratosphere) belt of very cold air, analogous to the one normally found above the intertropical convergence near the Equator, occurs above the anticyclonic ridge, across southern Asia at 30°–40° N and above the 500-millibar level (6,000 metres [19,700 feet]). These upper-air features that arise so far away from the Equator are associated with the surface monsoon and are absent when there is no monsoonal flow. The position of the easterly jet controls the location of monsoonal rains, which occur ahead and to the left of the strongest winds and also behind them and to the right. The surface flow, however, is a strong, southwesterly, humid, and unstable wind that brings humidities of more than 80 percent and heavy, squally showers that are the “burst” of the monsoon. The overall pattern of the advance follows a frontal alignment (front), but local episodes may differ considerably. The amount of rain is variable from year to year and place to place.

      Most spectacular clouds and rain occur against the Western Ghats (India) in India, where the early monsoonal airstream piles up against the steep slopes, then recedes, and piles up again to a greater height. Each time it pushes thicker clouds upward until wind and clouds roll over the barrier and, after a few brief spells of absorption by the dry inland air, cascade toward the interior. The windward slopes receive from 2,000 to 5,000 mm (80 to 200 inches) of rain in the monsoon season.

      Various factors, especially topography, combine to make up a complex regional pattern. Oceanic air flowing toward India below 6,000 metres (19,700 feet) is deflected in accordance with the Coriolis effect (Coriolis force) (see Relationship of wind to pressure and governing forces (climate)). The converging, moist oncoming stream becomes unstable over the hot land and is subject to rapid convection. Towering cumulonimbus clouds rise thousands of metres, producing violent thunderstorms (thunderstorm) and releasing latent heat in the surrounding air. As a result, the upper tropospheric warm belt migrates northwestward from the ocean to the land. The main body of air above 9,000 metres (29,500 feet) maintains a strong easterly flow.

      Later, in June and July, the monsoon is strong and well-established to a height of 6,000 metres (less in the far north), with occasional thickening to 9,000 metres. weather conditions are cloudy, warm, and moist all over India. Rainfall varies between 400 and 500 mm (16 and 20 inches), but topography introduces some extraordinary differences. On the southern slopes of the Khasi Hills at only 1,300 metres (4,300 feet), where the moist airstreams are lifted and overturned, the village of Cherrapunji in Meghalaya (Meghālaya) state receives an average rainfall of 2,730 mm (107 inches) in July, with record totals of 897 mm (35 inches) in 24 hours in July 1915, more than 9,000 mm (354 inches) in July 1861, and 16,305 mm (642 inches) in the monsoon season of 1899. Over the Ganges (Ganges River) valley the monsoon, deflected by the Himalayan (Himalayas) barrier, becomes a southeasterly airflow. By then the upper tropospheric belt of warmth from condensation has moved above northern India, with an oblique bias. The lowest pressures prevail at the surface.

      It is mainly in July and August that waves of low pressure appear in the body of monsoonal air. Fully developed depressions appear once or twice per month. They travel from east to west more or less concurrently with high-level easterly waves and bursts of speed from the easterly jet, causing a local strengthening of the low-level monsoonal flow. The rainfall consequently increases and is much more evenly distributed than it was in June. Some of the deeper depressions become tropical cyclones (tropical cyclone) before they reach the land, and these bring torrential rains and disastrous floods (flood).

      A totally different development arises when the easterly jet moves farther north than usual. The monsoonal wind rising over the southern slopes of the Himalayas brings heavy rains and local floods. The weather over the central and southern districts, however, becomes suddenly drier and remains so for as long as the abnormal shift lasts. The opposite shift is also possible, with midlatitude upper air flowing along the south face of the Himalayas and bringing drought to the northern districts. Such dry spells are known as “breaks” of the monsoon. Those affecting the south of India are similar to those experienced on the Guinea Coast during extreme northward shifts of the wind belts (as later discussed in the section on the West African monsoon (climate)), whereas those affecting the north are due to an interaction of the middle and low latitudes. The southwest monsoon over the lower Indus (Indus River) Plain is only 500 metres (about 1,600 feet) thick and does not hold enough moisture to bring rain. On the other hand, the upper tropospheric easterlies become stronger and constitute a true easterly jet stream. Western Pakistan, Iran, and Arabia remain dry (probably because of the divergence in this jet) and thus become the new source of surface heat.

      By August the intensity and duration of sunshine (sunlight) have decreased, temperatures begin to fall, and the surge of southwesterly air diminishes spasmodically almost to a standstill in the northwest. Cherrapunji still receives over 2,000 mm (79 inches) of rainfall at this time, however. In September, dry, cool, northerly air begins to circle the west side of the highlands and spread over northwestern India. The easterly jet weakens, and the upper tropospheric easterlies move much farther south. Because the moist southwesterlies at lower levels are much weaker and variable, they are soon pushed back. The rainfall becomes extremely variable over most of the region, but showers are still frequent in the southeastern areas and over the Bay of Bengal.

      By early October, variable winds are very frequent everywhere. At the end of the month, the entire Indian region is covered by northerly air and the winter monsoon takes shape. The surface flow is deflected by the Coriolis force and becomes a northeasterly flow. This causes an October–December rainy season for the extreme southeast of the Deccan (including the Madras coast) and eastern Sri Lanka, which cannot be explained by topography alone because it extends well out over the sea. Tropical depressions and cyclones are important contributing factors.

 Most of India thus begins a sunny, dry, and dusty season. The driest period comes in November in the Punjab; December in central India, Bengal, and Assam; January in the northern Deccan; and February in the southern Deccan. Conversely, the western slopes of the Karakoram Range and Himalayas are then reached by the midlatitude frontal depressions that come from the Atlantic and the Mediterranean. The winter rains they receive, moderate as they are, place them clearly outside the monsoonal realm.

      Because crops and water supplies depend entirely on monsoonal rains, it became imperative that quantitative, long-range weather forecasts be available. Embedded in the weather patterns of other parts of the world are clues to the summer conditions in South Asia. These clues often appear in the months leading up to monsoon onset. For a forecast to be released at the beginning of June, South American pressure and Indian upper-wind data for the month of April are examined. These data, though widely separated from one another, are positively correlated and may be used as predictors of June conditions. Forecasts may be further refined in May by comparing rainfall patterns in both Zimbabwe and Java to the easterly winds above the city of Kolkata ( Calcutta) in West Bengal state. In this situation the correlation between rainfall and easterly winds is negative.

Air pollution and the winter monsoon in Asia
      During the winter months in Asia, the Siberian anticyclone, which is typically centred near Lake Baikal (Baikal, Lake) in Russia, extends its influence over most of central and northern China. In these locations, subsiding air and strong temperature inversions (temperature inversion) prevail in the lowest kilometre of the atmosphere. Temperature inversions create regions of stagnant air that have the effect of trapping air pollution (pollution). In China, Southeast Asia, and South Asia, the widespread practice of burning firewood and coal for cooking, heating, and industry produces significant air pollution in the lower troposphere. The prevailing northeasterly winds of the winter monsoon transport the bulk of this polluted air southward, where it contributes to what has become known as the “Asian Brown Cloud” over the Indian Ocean.

      A strong reversal of winds, where northeasterly winds near Earth's surface shift to more westerly directions above 1 km (0.6 mile), is a characteristic feature of the northeast monsoon. Parcels of polluted air rising above the 1-km level are carried eastward, over the Pacific Ocean where they disperse. When these parcels are caught in the strong winds of the subtropical jet stream, they have been known to traverse half the globe in a matter of 10 to 15 days. As a result, pollution related to the winter monsoon has become widely recognized as a major scientific issue with global environmental implications.

The Malaysian (Malaysia)-Australian (Australia) monsoon
      Southeast Asia and northern Australia are combined in one monsoonal system that differs from others because of the peculiar and somewhat symmetrical distribution of landmasses on both sides of the Equator. In this respect, the northwest monsoon of Australia is unique. The substantial masses of water between Asia and Australia have a moderating effect on tropospheric temperatures, weakening the summer monsoon. The many islands (e.g., Philippines and Indonesia) provide an infinite variety of topographic effects. Typhoons (typhoon) that develop within the monsoonal air bring additional complications.

 It would be possible to exclude North China, Korea, and Japan from the monsoonal domain because their seasonal rhythm follows the normal midlatitude pattern—a predominant outflow of cold continental air in winter and frontal depressions and rain alternating with fine, dry anticyclonic weather in the warm season. On the other hand, the seasonal reversal of wind direction in this area is almost as persistent as that in India. The winter winds of northeastern Asia are much stronger because of the relative proximity of the Siberian anticyclone. The tropical ridge of high pressure is the natural boundary between these non-monsoonal areas and the monsoonal lands farther south.

      The northern limit of the typical monsoon may be set at about 25° N latitude. Farther north the summer monsoon is not strong enough to overcome the effect of the traveling anticyclones normally typical of the subtropics. As a result, monsoonal rains occur in June and also in late August and September, separated by a mild anticyclonic drought in July. In South China and the Philippines the trade winds (trade wind) prevail in the October–April (winter) period, strengthened by the regional, often gusty, outflow of air from the stationary Siberian anticyclone. Their disappearance and replacement by opposite (southwesterly) winds in the May–September (summer) period is the essence of the monsoon. In any case, these monsoonal streams are quite shallow, about 1,500 metres (4,900 feet) in winter and 2,000 metres (about 6,600 feet) in summer. They bring rain only when subject to considerable cooling, such as anywhere along the steep, windward slopes of the Philippines and Taiwan. On the larger islands there are contrasting effects: the slopes facing west receive most of their rainfall from May to October and experience drought from December to April, whereas the slopes facing east receive orographic rains (those produced when moist air is forced to rise by topography) from September to April and mainly convectional rains from May to October.

      In Vietnam and Thailand the summer monsoon is more strongly developed because of the wider expanses of overheated land. The southwesterly stream flows from May to October, reaching a thickness of 4 to 5 km (about 2.5 to 3 miles); it brings plentiful, but not extraordinary, rainfall. The period from November through February is the cool, dry season, and the period from March through April is the hot, dry one; in the far south the coolness is but relative. Along the east coast and on the eastward slopes, more rain is brought by the winter monsoon. In the summer, somewhere between Thailand and Cambodia in the interior, there may be a faint line of convergence between the southwesterly Indian-Myanmar monsoon and the southeasterly Malaysian monsoon.

      Monsoonal winds are weak over Indonesia because of the expanses of water and the low latitude, but their seasonal reversal is definite. From April to October the Australian southeasterly air flows, whereas north of the Equator the flow becomes a southwesterly. The Malaysian-Australian monsoon generally maintains its dryness over the islands closer to Australia, but farther north it carries increasing amounts of moisture. The northeasterly flow from Asia, which becomes northwesterly south of the Equator, is laden with moisture when it reaches Indonesia, bringing cloudy and rainy weather between November and May. The wettest months are December in most of Sumatra and January elsewhere, but rainfall patterns are highly localized. In Java, for instance, at sea level alone there are two major regions: an “equatorial” west with no dry season and a “monsoonal” east with extreme drought in August and September.

      Because of its relatively small size and compact shape, Australia shows relatively simple monsoonal patterns. The north shore is subject to a clear-cut wind reversal between summer (November–April, northwesterly flow) and winter (May–September, southeasterly flow) but with two definite limitations: first, the northwesterly, rain-bearing monsoonal wind is often held offshore and is most likely to override the land to any depth during January and February; second, even in summer there often are prolonged spells of southeasterly trade winds issuing from traveling anticyclones, separating the brief monsoonal incursions. The Australian summer monsoon is thus typical in direction and weather type but quite imperfect in frequency and persistence. Its thickness is usually less than 1,500 metres (4,900 feet) over the sea and 2,000–2,500 metres (6,600–8,200 feet) over the land.

      Much less typical are the marginal monsoonal manifestations. On the northwest coast there frequently is a northwesterly airflow in the summer (December–March), as opposed to the winter southeasterlies, but this stream is very shallow and does not bring any rain—i.e., its weather is not monsoonal even though its direction is so. On the northeast coast the onshore air is humid and brings rain, but its direction is only partly modified in summer. Most of the summer winds that arrive here occur as a northeasterly flow, although at other times the flow can be mostly southeasterly.

The West African monsoon
 The main characteristics of the West African (Western Africa) seasons have been known to the scientific community for more than two centuries. The southwest winter monsoon flows as a shallow (less than 2,000 metres [about 6,600 feet]) humid layer of surface air overlain by the primary northeast trade wind, which blows from the Sahara and the Sahel as a deep stream of dry, often dusty air. As a surface northeasterly, it is generally known as the harmattan, gusty and dry in the extreme—cool at night and scorchingly hot by day. As in a thorough monsoonal development, upper tropospheric anticyclones occur at about 20° N, while the easterly jet stream may occur at about 10° N, much closer to the Equator than they are in the Indian region.

 The West African monsoon is the alternation of the southwesterly wind and the harmattan at the surface. Such alternation is normally found between latitudes 9° and 20° N. Northeasterlies occur constantly farther north, but only southwesterlies occur farther south. Except for erratic rains in the high-sun season, the whole year is more or less dry at 20° N. The drought becomes shorter and less complete farther south, as shown in the map—>. At 12° N it lasts about half the year, and at 8° N it disappears completely. Farther south a different, lighter drought begins to appear in the high-sun months when the monsoonal southwesterly is strongest. This drought is due to the arrival of dry surface air issuing from anticyclones formed beyond the Equator in the Southern Hemisphere and is thus similar to the monsoonal drought in Java. Like the “break” of the monsoon in southern India, however, it occurs beyond the Equator.

      The moist southwesterly stream, particularly frequent between 5° and 10° N, can reach much farther north, bringing warm, humid nights and moderately hot, but still humid, days. The harmattan brings cooler nights, but the extreme daily heating causes a thermal range of 10–12 °C (18–22 °F). Even in the daytime, the harmattan may give a sensation of coolness to the human skin as it evaporates moisture from the skin's surface. The alternation of the two winds is seasonal on the basis of overall frequency, but in fact it varies considerably with the synoptic pressure patterns. The harmattan comes in spells that mostly last from a few days to more than a week.

      The advancing fringe of the southwest monsoon is too shallow (under 1,000 metres [3,300 feet]) for many thunderstorms and other disturbances to occur. They usually occur 200–300 km (about 125–185 miles) behind the fringe, where the moist air is deeper (1,000–2,000 metres [3,300–6,600 feet]) but the ground is still hot enough to make it very unstable. The tops of cumulonimbus clouds may reach 12,000 metres (about 39,000 feet), well above freezing level (4,200–4,500 metres [13,800–14,800 feet]). The disturbances usually occur along a given longitude line that is slightly curved and may in fact form one long line squall. They also reach 12,000 metres or more, traveling steadily westward at 37 to 56 km (23 to 35 miles) per hour. This suggests that they originate in the primary trade wind aloft and, as in India, are probably related to the tropical easterly jet stream. The southwest monsoon dominates the weather, and clouds and rain abound. The rain is primarily due to coalescence of droplets, with most of the clouds located below 3,500 metres (11,500 feet). The humidity is very high, and the daily range of temperature remains around 4 °C (7 °F).

      If it were not for the change in wind direction when the southeast trades have crossed the Equator, the monsoon system of West Africa could not be distinguished from the weather system, caused by the seasonal shift in the latitude of the intertropical convergence, as experienced over most of Central Africa. There is a rainy season (in this case, the monsoon season), which lasts two to three months at latitude 16° N on the west coast, three to four months at 14° N, six to seven months at 10° N, and eight months at 8° N.

      On the south coast, which is at latitude 4° N to 6° N, the southwest monsoon (as the intertropical convergence) may occur at any time, but the results are quite atypical for various reasons. In the low-sun season (December–February), the southwesterly is rare and ineffective, and the weather is cloudy but dry. From April to June, the midday sun is at its highest, and insolation (radiation received at Earth's surface) is most intense. Because the southwest wind occurs most frequently, the consequent building up of clouds leads to the main rainy season. During July and August (the short drought), cloudy conditions prevail, but the air issues direct from anticyclones farther south and is dry, in spite of the fact that its direction of flow does not change. Although cloudiness decreases after the second high-sun season in September and October, there is a period of occasional rains just sufficient to constitute a secondary maximum.

      Toward the north, conditions are more distinctly monsoonal: by latitude 8° N the two wet seasons have merged into one long “wet” with two subdued peaks, which last approximately seven to eight months (March–October). The “dry,” which is controlled by northeast winds, lasts from November to early March. There is one rainfall maximum (in August or September) only a short distance farther north, although the wet season is only a few weeks shorter.

Monsoonal tendencies in Europe and North America
 In central Europe, where the average wind direction in summer differs some 30° to 40° from that of the Atlantic, there are monsoonal tendencies that occur not as a continuous flow but rather intermittently within frontal depressions, bringing cool, cloudy weather, rain, and thunderstorms. Some see in this climatic pattern a true monsoon, but it is obvious that it is only an “embryo monsoon” that results in weather singularities. The latitude is too high for a true monsoon to arise.

      In North America the relatively low latitude and the orientation of the land-sea boundary on the Gulf of Mexico (Mexico, Gulf of) are quite favourable to monsoonal developments. During the summer, low atmospheric pressure is frequent over the heated land; the northeasterly trade winds are consequently deflected to become easterly, southeasterly, or even southerly winds. In general, Texas and the Gulf Coast of the United States may be completely overrun by a shallow sheet of oceanic air, which may continue for a long distance inland. The rainfall regime does not reveal any marked monsoonal pattern. There are mostly two, three, or even four minor peaks in the sequence of monthly rainfall totals. In the winter there often occur “northers,” which are offshore winds caused by the general anticyclonic flow of air from the cold land. Neither the summer onshore wind nor the winter offshore wind is persistent enough to constitute a monsoonal sequence, even though monsoonal tendencies are quite evident.

      In Central America, a true monsoonal cycle occurs over a small area facing the Pacific Ocean between 5° and 12° N. Not only is there a complete seasonal reversal of the wind, but the rainfall regime is typically monsoonal. The winter period (from November to January and from March to April according to latitude and other factors) is very dry. The rainy season begins earlier (May) in the south and progressively later farther north, coming at the end of June in southern Mexico. It concludes at the end of September in the north and as late as early November in the south. The result is a rainy season that increases in duration with decreasing latitude; it lasts three months in southern Mexico and from six to seven months in Costa Rica. Latitude for latitude, this is a subdued replica of the monsoon of India.

Monsoon variability

Diurnal variability
      Landmasses in regions affected by monsoons warm up very rapidly in the afternoon hours, especially on days with cloud-free conditions; surface air temperatures between 35 and 40 °C (95 and 104 °F) are not uncommon. Under such conditions, warm air is slowly and continually steeped in the moist and cloudy environment of the monsoon. Consequently, over the course of a 24-hour period, energy from this pronounced diurnal, or daily, change in terrestrial heating is transferred to the cloud, rain, and diurnal circulation systems. The scale of this diurnal change extends from that of coastal sea breezes to continent-sized processes. Satellite observations have confirmed that the effects of rapid diurnal temperature change occur at continental scales. For example, air from surrounding areas is drawn into the lower troposphere over warmer land areas of South Asia during summer afternoon hours. This buildup of afternoon heating is accompanied by the production of clouds and rain. In contrast, a reverse circulation, characterized by suppressed clouds and rain, is noted in the early morning hours.

Intra-annual variability
      Monsoon rainfall and dry spells alternate on several timescales. One such well-known timescale is found around periods of 40–50 or 30–60 days. This is called the Madden-Julian oscillation (MJO), named for American atmospheric scientists Roland Madden and Paul Julian in 1971. This phenomenon comes in the form of alternating cyclonic and anticyclonic regions that enhance and suppress rainfall, respectively, and flow eastward along the Equator in the Indian and Pacific oceans. The MJO has the ability to influence monsoonal circulation and rainfall by adding moisture during its cyclonic (wet) phase and reducing convection during its anticyclonic (dry) phase. At the surface in monsoon regions, both dry and wet spells result. These periods may alternate locally on the order of two or more weeks per phase.

Interannual variability
 The variability of monsoon-driven rainfall in the Indian Ocean and Australia appears to parallel El Niño episodes. During El Niño events, which occur about every two to seven years, ocean temperatures rise over the central equatorial Pacific Ocean by about 3 °C (5.4 °F). Atypical conditions characterized by increased rising air motion, convection, and rain are created in the western equatorial Pacific. At the same time, a compensating lobe of descending air, producing below-normal rainfall, appears in the vicinity of eastern Australia, Malaysia, and India. The graph illustrates a well-known El Niño–monsoon rainfall relationship. Here, precipitation figures from above- and below-normal monsoon rainfall periods over India are expressed as a function of years. Years characterized by El Niño events are marked by darkened histogram barbs. The graph shows that many of the years with below-normal monsoon rainfall coincide with El Niño years. This illustration provides only limited guidance to seasonal forecasters since monsoon rainfall is close to normal during many El Niño and La Niña years.

      Many other factors, aside from equatorial Pacific Ocean surface temperatures, contribute to the interannual variability of monsoon rainfall. Excessive spring snow and ice cover on the Plateau of Tibet is related to the deficient monsoon rainfall that occurs during the following summer season in India. Furthermore, strong evidence exists that relates excessive snow and ice cover in western Siberia to deficient Asian summer rainfall. Warmer than normal sea surface temperatures over the Indian Ocean may also contribute somewhat to above-normal rainfall in South Asia. The interplay among these many factors makes forecasting monsoon strength a challenging problem for researchers.

 A rather clear signature on the decadal variability of Indian rainfall has been documented by the Indian Weather Service. Decadal-scale variability appears in the graph as an annual running mean that combines average rainfall anomalies (totals as a departure from normal rainfall amounts) occurring at all Indian rain gauge sites. Periods of heavier-than-normal rainfall are followed by decades of somewhat less rainfall.

Joseph Gentilli Phillip J. Smith T.N. Krishnamurti

Upper-level winds
 The flow of air around the globe is greatest in the higher altitudes, or upper levels. Upper-level airflow occurs in wavelike currents that may exist for several days before dissipating. Upper-level wind speeds generally occur on the order of tens of metres per second and vary with height. The characteristics of upper-level wind systems vary according to season and latitude and to some extent hemisphere and year. Wind speeds are strongest in the midlatitudes near the tropopause and in the mesosphere.

      Upper-level wind systems, like all wind systems, may be thought of as having parts consisting of uniform flow, rotational flow (with cyclonic (cyclone) or anticyclonic (anticyclone) curvature), convergent or divergent flow (convergence and divergence) (in which the horizontal area of masses of air shrinks or expands), and deformation (by which the horizontal area of air masses (air mass) remains constant while experiencing a change in shape). Upper-level wind systems in the midlatitudes tend to have a strong component of uniform flow from west to east (“westerly (Pacific Ocean)” flow), though this flow may change during the summer. A series of cyclonic and anticyclonic vortices superimposed on the uniform west-to-east flow make up a wavetrain (a succession of waves occurring at periodic intervals). The waves are called Rossby waves (Rossby wave) after Swedish American meteorologist C.G. Rossby (Rossby, Carl-Gustaf Arvid), who first explained fundamental aspects of their behaviour in the 1930s. Waves whose wavelengths (wavelength) are around 6,000 km (3,700 miles) or less are called “short waves,” while those with longer wavelengths are called “long waves.” In addition, short waves progress in the same direction as the mean airflow, which is from west to east in the midlatitudes; long waves retrogress (that is, move in the opposite direction of the mean flow). Although the undulating current of air is composed of a number of waves of varying wavelength, the dominant wavelength is usually around several thousand kilometres. Near and underneath the tropopause (troposphere), regions of divergence are found over regions of gently rising air at the surface, while regions of convergence aloft are found over regions of sinking air below. These regions are usually much more difficult to detect than the regions of rotational and uniform flow. While the horizontal wind speed is typically in the range of 10–50 metres per second (about 20–110 miles per hour), the vertical wind speed associated with the waves is only on the order of centimetres per second.

      The characteristics of upper-level wind systems are known mainly from an operational worldwide network of rawinsonde observations. (A rawinsonde is a type of radiosonde designed to track upper-level winds and whose position can be tracked by radar.) Winds measured from Doppler (Doppler effect)-radar wind profilers, aircraft navigational systems, and sequences of satellite-observed cloud imagery have also been used to augment data from the rawinsonde network; the latter two have been especially useful for defining the wind field over data-sparse regions, such as over the oceans.

      The winds at upper levels, where surface friction does not occur, tend to be approximately geostrophic (geostrophic motion). In other words, there is a near balance between the pressure (atmospheric pressure) gradient force, which directs air from areas of relatively high pressure to areas of relatively low pressure, and the Coriolis force, which deflects air from its straight-line path to the right in the Northern Hemisphere and to the left in the Southern Hemisphere. An important consequence of this geostrophic balance is that the winds blow parallel to isobars (cartographic lines indicating areas of equal pressure), and, according to Buys Ballott's law (Buys Ballot's law), lower pressures will be found to the left of the direction of the wind in the Northern Hemisphere and to the right of the wind in the Southern Hemisphere. Furthermore, wind speed increases as the spacing between isobars (isobar) decreases. In a wavetrain of westerly flow, the regions of cyclonic flow are associated with troughs of low pressure, whereas anticyclonic flow are characterized by ridges of high pressure. Rising motions tend to be found downstream from the troughs and upstream from the ridges, while sinking motions tend to be found downstream from the ridges and upstream from the troughs. The areas of rising motion tend to be associated with clouds (cloud) and precipitation (inclement weather), whereas the areas of sinking motion tend to be associated with clear skies (fair weather).

      The vertical variation of the structure of the waves depends upon the temperature pattern. In general, owing to the net difference in incoming shorter-wavelength solar radiation and outgoing longer-wavelength infrared radiation between the polar and the equatorial regions, there is a horizontal temperature gradient in the troposphere. At both the surface and upper levels, the troposphere is warmest at low latitudes and coldest at high latitudes. The atmosphere is mainly in hydrostatic balance, or equilibrium, between the upward-directed pressure gradient force and the downward-directed force of gravity. This circumstance is expressed in the following relationship:

p/∂z = –ρg (1)
where ∂p/∂z is the partial derivative of p with respect to z, p is the pressure, z is the height, ρ is the density of the air, and g is the acceleration of gravity. A consequence of this hydrostatic relationship is that the pressure at any level is equal to the weight of the column of air above. According to the ideal gas law,
p = ρRT (2)
where R is the gas constant and T is the temperature. At any given pressure, the density varies inversely with temperature. Therefore, relatively cold air is heavier than relatively warm air at the same pressure. It follows from (1) and (2) that pressure decreases more rapidly with height at high latitudes in the colder air than it does at lower latitudes in the warmer air. If there is a westerly geostrophic wind at midlevels in the troposphere, then pressure decreases with increasing latitude. Consequently, the horizontal spacing between isobars decreases with height. Thus, the geostrophic wind speed, which approximates the actual wind speed, increases with height. Above the tropopause the pole-to-Equator temperature gradient is reversed as air temperature increases with height, so that the westerlies decrease in intensity in the stratosphere. Thus, the strongest westerly current of winds is located near the tropopause.

      The aforementioned relationship can be analyzed quantitatively by considering the vertical variation in the geostrophic wind, which is found from the hydrostatic equation (1), the ideal gas law (perfect gas) (2), and the geostrophic wind formula, approximately as follows.

ug/∂z = – g/fTT/∂y and ∂vg/∂z = – g/fTT/∂z, (3)
where ∂ug/∂z is the partial derivative of ug with respect to z, ug and vg are the components of the geostrophic wind in the zonal (straight from west to east) and meridional (north to south) directions, respectively, and f is the Coriolis parameter. The equations given in (3) are known as the thermal-wind relations. The difference between the geostrophic wind at some higher level and the geostrophic wind below is called the thermal wind. It follows that the thermal wind vector is oriented so that in the colder air it lies to the left in the Northern Hemisphere and to the right in the Southern Hemisphere.

      In addition to the general pole-to-Equator temperature gradient found in the troposphere, there are zonally oriented temperature variations that are wavelike. In fact, to a first approximation, the isotherms (isotherm) (cartographic lines indicating areas of equal temperature) are nearly parallel to the isobars in the upper levels of the troposphere. Most frequently, relatively cold air lies just upstream from upper-level troughs and just downstream from upper-level ridges, while relatively warm air lies just upstream from upper-level ridges and just downstream from upper-level troughs. The thermal-wind relation (3) indicates that the wavetrain of troughs and ridges tilts with height to the west. In the midlatitudes during the summer, and in some locations within the midlatitudes during the winter, the meridional temperature gradient weakens so much that the westerlies become weak or nonexistent. As a result, the wavelike wind field disappears and the flow pattern is that of cyclones and anticyclones “cut off” from the flow. When cold air is colocated with the upper-level cyclones and warm air is colocated with the upper-level anticyclones, according to (3), both circulation patterns increase in intensity with height and are called cold-core and warm-core systems, respectively. Tropical cyclones (tropical cyclone), on the other hand, are warm-core systems that are most intense at the surface and that decrease in intensity with height.

      The vertical structure of upper-level waves has an important effect on smaller-scale features that may be embedded within them. The susceptibility of the atmosphere to vertical overturning (a mixing of lower-level warmer air with higher-level colder air) through deep cumulus convection (e.g., thunderstorms (thunderstorm)) depends on the rate at which temperature decreases with height. When regions of relatively cold air aloft associated with upper-level troughs or cyclones become superimposed during the winter over relatively warm ocean surfaces or during the summer over hot and humid landmasses, then convective storms can form. The type of mesoscale convective system (MCS) that can form depends in large part on the vertical wind shear. When the vertical shear is very strong, supercells and tornadoes (tornado) may be spawned, especially during the warmer months. During the winter, bands of precipitation sometimes line up along the vertical shear vector through a process known as slantwise convection.

Propagation and development of waves
      Upper-level waves in the westerlies in midlatitudes usually move from west to east, in part as a result of advection (a process in which the airflow transports a property of the atmosphere [warmth, cold, etc.] downstream) and in part as a result of propagation, which acts in the opposite direction, toward the west. Rossby showed that to a good approximation,

c = U – β / (2π/L)2, (4)
where c is the phase speed of the waves, U is the speed from west to east of the component of upper-level wind due to uniform flow, β is the meridional, or north-south, gradient of the Coriolis parameter (f), and L is the zonal wavelength (the distance between successive troughs or ridges). According to (4), since the magnitude of f increases toward the poles, β is positive, and hence waves whose wavelengths are short have a relatively small component due to propagation. In this situation, advection overwhelms the effect of propagation and the waves move on downstream. On the other hand, if in midlatitudes the wavelength is very long, then the effects of propagation may exactly cancel the effects of advection, and the waves may become stationary; or if the wavelength becomes even longer, then the waves may become retrograde. From (4) it can be seen that Rossby waves owe their propagation characteristics to the north-south variation of f. In nature, temperature effects and heating and cooling over warm and cold surfaces can modify (4) somewhat.

      The physical basis for (4) and for the development of upper-level systems and how they relate to surface systems is described by an elegant theory developed in the late 1940s called quasigeostrophic theory. A measure of the tendency for a fluid to rotate is known as vorticity and is given by the following equation:

ζ = ∂v/∂x – ∂u/∂y (5)
where ζ is the relative vorticity with respect to Earth's surface. The variables x and y are the coordinate axes for space and correspond to the measurements to the east and north, respectively. The variables u and v are zonal and meridional components (the components of motion in the easterly and northerly directions), respectively, of the wind. On the rotating Earth, the vorticity is the sum of the relative vorticity with respect to Earth's surface, given by the aforementioned expression, and Earth's vorticity, given by f, the Coriolis parameter. Troughs are associated with cyclonic vorticity, and ridges are associated with anticyclonic vorticity. In a wavetrain, the pressure falls downstream from troughs, where the wind is directed from the region of maximum vorticity along the trough to the region of minimum vorticity, which is along the ridge, and the pressure rises downstream from ridges. On the other hand, pressure can rise east of troughs (and west of ridges) where there is component of motion from the Equator to the pole. For example, pressure rises from regions of low magnitude of f to higher magnitude of f (from low values of Earth's vorticity to higher values of Earth's vorticity). Likewise, pressure falls west of troughs (and east of ridges) where there is a component of motion from one of the poles to the Equator—from relatively high magnitude of f to lower magnitude of f. The effect of pressure increases and decreases is greatest when the wavelength is relatively short, such as when the effects of the advection of Earth's vorticity are overwhelmed by the effects of advection of relative vorticity.

      The development and amplification of Rossby waves is typically a result of the advection of warmer or colder air at low levels. When warm air is advected underneath a layer of air not experiencing much, if any, advection, the pressure at the top of the layer rises. Conversely, pressure falls when cold air advects under a similar layer of air. If the wavetrain tilts to the west with height so that cold air lies to the west of troughs and thus east of ridges, the pressure aloft in the troughs decreases. Similarly, when warm air lies to the east of troughs and thus west of ridges, the pressure aloft in the ridges increases. As a result, the amplitude of the waves in the wavetrain increases, thereby enhancing the temperature advection process, so that there is a positive feedback mechanism that makes the waves continue to amplify. In this process, called “baroclinic instability,” potential energy is converted into kinetic energy—which occurs as wind—as warm, light air rises and cold, heavy air sinks. Since baroclinic instability is associated with horizontal temperature gradients, according to the thermal wind relation (3), there must be vertical wind shear.

      It is also possible for Rossby waves to amplify through a process called barotropic instability. Barotropic instability, however, requires horizontal shear, not vertical shear; kinetic energy for the waves comes from the mean kinetic energy associated with the westerly wind current. The waves grow in amplitude at the expense of the mean flow. Barotropic instability can occur when the horizontal shear varies with latitude such that the sum of Earth's vorticity and the relative vorticity associated with the horizontal shear is small with respect to latitude.

Relationships to surface features
 Rossby waves (Rossby wave) propagating through the upper and middle troposphere cause disturbances to form at the surface. According to quasigeostrophic theory, when there is a wavetrain embedded within a zone of pole-to-Equator temperature gradient, air rises east of upper-level troughs (and west of upper-level ridges) and sinks west of upper-level troughs (and east of upper-level ridges). These vertical air motions are required to maintain the approximate geostrophic and hydrostatic balance, which are necessary for quasigeostrophic equilibrium. Air converges at the surface underneath the rising current of air to compensate for the upward loss of mass and diverges at the surface underneath a sinking current of air to compensate for the downward gain of mass. As a consequence of the lateral deviation of the air by the Coriolis force, Earth's vorticity is converted into cyclonic relative vorticity where air converges and anticyclonic relative vorticity where air diverges. According to the geostrophic wind relation, cyclonic gyres are associated with low-pressure centres, whereas anticyclonic gyres are connected with areas of high pressure. Thus, low-pressure areas form at the surface downstream from upper-level troughs and upstream from upper-level ridges, whereas the reverse is true for high-pressure areas. These surface low- and high-pressure areas thereby create a westward tilt with height of the waves in pressure. Since there tends to be a pole-to-Equator-directed geostrophic wind west of surface lows and east of surface highs, and an Equator-to-pole-directed geostrophic wind east of surface lows and west of surface highs, there is cold advection underneath upper-level troughs and warm advection underneath upper-level ridges; the baroclinic instability process is thus facilitated.

Jet streams (jet stream)
 The upper-level wind flow described above is frequently concentrated into relatively narrow bands called jet streams (jet stream), or jets. The jets, whose wind speeds are usually in excess of 30 metres per second (about 70 miles per hour) but can be as high as 107 metres per second (about 240 miles per hour), act to steer upper-level waves. Jet streams are of great importance to air travel because they affect the ground speed, the velocity relative to the ground, of aircraft. Since strong upper-level flow is usually associated with strong vertical wind shear, jet streams in midlatitudes are accompanied by strong horizontal temperature gradients, as required by the thermal wind relation (3). Some regions of high vertical wind shear are marked by clear-air turbulence (CAT). Jet streams whose extents are relatively isolated are called jet streaks. Well-defined circulation patterns of rising and sinking air are usually found just upstream and downstream, respectively, from jet streaks (that are not too curved). Rising motion is found to the left and right just downstream and upstream, respectively, and sinking motion is found to the right and left just downstream and upstream, respectively. Jets tend to be strongest near the tropopause where the horizontal temperature gradient reverses.

      The polar front jet moves in a generally westerly direction in midlatitudes, and its vertical wind shear which extends below its core is associated with horizontal temperature gradients that extend to the surface. As a consequence, this jet manifests itself as a front that marks the division between colder air over a deep layer and warmer air over a deep layer. The polar front jet can be baroclinically unstable and break up into waves. The subtropical jet is found at lower latitudes and at slightly higher elevation, owing to the increase in height of the tropopause at lower latitudes. The associated horizontal temperature gradients of the subtropical jet do not extend to the surface, so that a surface front is not evident. In the tropics an easterly jet is sometimes found at upper levels, especially when a landmass is located poleward of an ocean, so the temperature increases with latitude. The polar front jet and the subtropical jet play a role in maintaining Earth's general circulation. They are slightly different in each hemisphere because of differences in the distribution of landmasses and oceans.

Winds in the stratosphere and mesosphere
 The winds in the stratosphere and mesosphere are usually estimated from temperature data collected by satellites. The winds at these high levels are assumed to be geostrophic. Overall, in the midlatitudes, they have a westerly component in the winter and an easterly component in the summer. The highest zonal winds are around 60–70 metres per second (135–155 miles per hour) at 65–70 km (37–43 miles) above Earth's surface. The west-wind component is stronger during the winter in the Southern Hemisphere. The axes of the strongest easterly and westerly wind components in the Southern Hemisphere tilt toward the south with increased altitude during the Northern Hemisphere winter and the Southern Hemisphere summer. The zonal component of the thermal wind shear is in accord with the zonal distribution of temperature.

      During the winter there is, in the mean, an intense cyclonic vortex about the poles in the lower stratosphere. Over the North Pole this vortex has an embedded mean trough over northeastern North America and over northeastern Asia, whereas over the Pacific there is a weak anticyclonic vortex. The winter cyclonic vortex over the South Pole is much more symmetrical than the one over the North Pole. During the summer there is an anticyclone above each pole that is much weaker than the wintertime cyclone.

      In the stratosphere, deviations from the mean behaviour of the winds occur during events called sudden warmings, when the meridional temperature gradient reverses on timescales as short as several days. This also has the effect of reversing the zonal wind direction. Sudden warmings tend to occur during the early and middle parts of the winter and the transition period from winter to spring. The latter marks the changeover from the cold winter polar cyclone to the warm summer polar anticyclone. It is noteworthy that long waves from the troposphere can propagate into the stratosphere during the winter when westerlies and sudden warmings occur, but this is not the case during the summer when easterly winds prevail.

      The zonal component of the winds in the stratosphere above equatorial and tropical regions is, in the mean, relatively weak. This is not necessarily the case at any given time, because they reverse direction on the average every 13–14 months. This phenomenon, which is known as the quasi-biennial oscillation (QBO), is caused by the interaction of vertically propagating waves with the mean flow. Its effect is greatest around 27 km (17 miles) above Earth's surface in the equatorial region. The strongest easterlies are stronger than the strongest westerlies.

Howard B. Bluestein

Climatic classification

General considerations
 The climate of an area is the synthesis of the environmental conditions (soils (soil), vegetation, weather, etc.) that have prevailed there over a long period of time. This synthesis involves both averages of the climatic elements and measurements of variability (such as extreme values and probabilities). Climate is a complex, abstract concept involving data on all aspects of Earth's environment. As such, no two localities on Earth may be said to have exactly the same climate. Nevertheless, it is readily apparent that, over restricted areas of the planet, climates vary within a limited range and that climatic regions are discernible within which some uniformity is apparent in the patterns of climatic elements. Moreover, widely separated areas of the world possess similar climates when the set of geographic relationships occurring in one area parallels that of another. This symmetry and organization of the climatic environment suggests an underlying worldwide regularity and order in the phenomena causing climate (e.g., patterns of incoming solar radiation, vegetation, soils, winds (wind), temperature, and air masses (air mass)), which were discussed in earlier sections.

      Climate classification is an attempt to formalize this process of recognizing climatic similarity, of organizing, simplifying, and clarifying the vast amount of environmental data, and of systematizing the long-term effects of interacting climatic processes to enhance scientific understanding of climates. Users of climate classifications should be aware of the limitations of the procedure, however.

      First, climate is a multidimensional concept, and it is not an obvious decision as to which of the many observed environmental variables should be selected as the basis of the classification. This choice must be made on a number of grounds, both practical and theoretical. For example, using too many different elements opens up the possibilities that the classification will have too many categories to be readily interpreted and that many of the categories will not correspond to real climates. Moreover, measurements of many of the elements of climate are not available for large areas of the world or have been collected for only a short time. The major exceptions are soil, vegetation, temperature, and precipitation data, which are more extensively available and have been recorded for extended periods of time.

      The choice of variables also is determined by the purpose of the classification (e.g., to account for distribution of natural vegetation, to explain soil formation processes, or to classify climates in terms of human comfort). The variables relevant in the classification will be determined by this purpose, as will the threshold values of the variables chosen to differentiate climatic zones.

      A second difficulty results from the generally gradual nature of changes in the climatic elements over Earth's surface. Except in unusual situations due to mountain ranges (mountain) or coastlines, temperature, precipitation, and other climatic variables tend to change only slowly over distance. As a result, climate types tend to change imperceptibly as one moves from one locale on Earth's surface to another. Choosing a set of criteria to distinguish one climatic type from another is thus equivalent to drawing a line on a map to distinguish the climatic region possessing one type from that having the other. While this is in no way different from many other classification decisions that one makes routinely in daily life, it must always be remembered that boundaries between adjacent climatic regions are placed somewhat arbitrarily through regions of continuous, gradual change and that the areas defined within these boundaries are far from homogeneous in terms of their climatic characteristics.

      Most classification schemes are intended for global- or continental-scale application and define regions that are major subdivisions of continents hundreds to thousands of kilometres across. These may be termed macroclimates. Not only will there be slow changes (from wet to dry, hot to cold, etc.) across such a region as a result of the geographic gradients of climatic elements over the continent of which the region is a part, but there will exist mesoclimates within these regions associated with climatic processes occurring at a scale of tens to hundreds of kilometres that are created by elevation differences, slope aspect, bodies of water, differences in vegetation cover, urban areas, and the like. Mesoclimates, in turn, may be resolved into numerous microclimates (microclimate), which occur at scales of less than 0.1 km (0.06 mile), as in the climatic differences between forests, crops, and bare soil, at various depths in a plant canopy, at different depths in the soil, on different sides of a building, and so on.

      These limitations notwithstanding, climate classification plays a key role as a means of generalizing the geographic distribution and interactions among climatic elements, of identifying mixes of climatic influences important to various climatically dependent phenomena, of stimulating the search to identify the controlling processes of climate, and, as an educational tool, to show some of the ways in which distant areas of the world are both different from and similar to one's own home region.

Approaches to climatic classification
      The earliest known climatic classifications were those of classical Greek (ancient Greek civilization) times. Such schemes generally divided the Earth into latitudinal zones based on the significant parallels of 0°, 23.5°, and 66.5° of latitude and on the length of day. Modern climate classification has its origins in the mid-19th century, with the first published maps of temperature and precipitation over Earth's surface, which permitted the development of methods of climate grouping that used both variables simultaneously.

      Many different schemes of classifying climate have been devised (more than 100), but all of them may be broadly differentiated as either empiric (Empiricism) or genetic methods. This distinction is based on the nature of the data used for classification. Empirical methods make use of observed environmental data, such as temperature, humidity, and precipitation, or simple quantities derived from them (e.g., evaporation (vaporization)). In contrast, a genetic method classifies climate on the basis of its causal elements, the activity and characteristics of all factors (air masses, circulation systems, fronts (front), jet streams (jet stream), solar radiation, topographic effects, and so forth) that give rise to the spatial and temporal patterns of climatic data. Hence, while empirical classifications are largely descriptive of climate, genetic methods are (or should be) explanatory. Unfortunately, genetic schemes, while scientifically more desirable, are inherently more difficult to implement because they do not use simple observations. As a result, such schemes are both less common and less successful overall. Moreover, the regions defined by the two types of classification schemes do not necessarily correspond; in particular, it is not uncommon for similar climatic forms resulting from different climatic processes to be grouped together by many common empirical schemes.

Genetic classifications
      Genetic classifications group climates by their causes. Among such methods, three types may be distinguished: (1) those based on the geographic (geography) determinants of climate, (2) those based on the surface energy budget (atmosphere), and (3) those derived from air-mass analysis.

      In the first class are a number of schemes (largely the work of German climatologists (climatology)) that categorize climates according to such factors as latitudinal control of temperature, continentality versus ocean-influenced factors, location with respect to pressure (atmospheric pressure) and wind belts, and effects of mountains. These classifications all share a common shortcoming: they are qualitative, so that climatic regions are designated in a subjective manner rather than as a result of the application of some rigorous differentiating formula.

      An interesting example of a method based on the energy balance of Earth's surface is the 1970 classification of Werner H. Terjung, an American geographer. His method utilizes data for more than 1,000 locations worldwide on the net radiation received at the surface, the available energy for evaporating water, and the available energy for heating the air and subsurface. The annual patterns are classified according to the maximum energy input, the annual range in input, the shape of the annual curve, and the number of months with negative magnitudes (energy deficits). The combination of characteristics for a location is represented by a label consisting of several letters with defined meanings, and regions having similar net radiation climates are mapped.

      Probably the most extensively used genetic systems, however, are those that employ air-mass concepts. Air masses (air mass) are large bodies of air that, in principle, possess relatively homogeneous properties of temperature, humidity, etc., in the horizontal. Weather on individual days may be interpreted in terms of these features and their contrasts at fronts.

      Two American geographer-climatologists have been most influential in classifications based on air mass. In 1951, Arthur N. Strahler described a qualitative classification based on the combination of air masses present at a given location throughout the year. Some years later (1968 and 1970), John E. Oliver placed this type of classification on a firmer footing by providing a quantitative framework that designated particular air masses and air mass combinations as “dominant,” “subdominant,” or “seasonal” at particular locations. He also provided a means of identifying air masses from diagrams of mean monthly temperature and precipitation plotted on a “thermohyet diagram,” a procedure that obviates the need for less common upper-air data to make the classification.

Empirical classifications
      Most empirical classifications are those that seek to group climates based on one or more aspects of the climate system. While many such phenomena have been used in this way, natural vegetation stands out as one of prime importance. The view held by many climatologists is that natural vegetation functions as a long-term integrator of the climate in a region; the vegetation, in effect, is an instrument for measuring climate in the same way that a thermometer measures temperature. Its preeminence is apparent in the fact that many textbooks and other sources refer to climates using the names of vegetation, as, for example, rainforest, taiga, or tundra.

      Wladimir Köppen (Köppen, Wladimir), a German botanist-climatologist, developed the most popular (but not the first) of these vegetation-based classifications. His aim was to devise formulas that would define climatic boundaries in such a way as to correspond to those of the vegetation zones that were being mapped for the first time during his lifetime. Köppen published his first scheme in 1900 and a revised version in 1918. He continued to revise his system of classification until his death in 1940. Other climatologists have modified portions of Köppen's procedure on the basis of their experience in various parts of the world.

      Köppen's classification is based on a subdivision of terrestrial climates into five major types, which are represented by the capital letters A, B, C, D, and E. Each of these climate types except for B is defined by temperature criteria. Type B designates climates in which the controlling factor on vegetation is dryness (rather than coldness). Aridity is not a matter of precipitation alone but is defined by the relationship between the precipitation input to the soil in which the plants grow and the evaporative losses. Since evaporation is difficult to evaluate and is not a conventional measurement at meteorological stations, Köppen was forced to substitute a formula that identifies aridity in terms of a temperature-precipitation index (i.e., evaporation is assumed to be controlled by temperature). Dry climates are divided into arid (BW) and semiarid (BS) subtypes, and each may be differentiated further by adding a third code, h for warm and k for cold.

       Classification of major climatic types according to the Koppen-Geiger-Pohl schemeAs noted above, temperature defines the other four major climate types. These are subdivided, with additional letters again used to designate the various subtypes. Type A climates (the warmest) are differentiated on the basis of the seasonality of precipitation: Af (no dry season), Am (short dry season), or Aw (winter dry season). Type E climates (the coldest) are conventionally separated into tundra (ET) and snow/ ice climates (EF). The mid-latitude C and D climates are given a second letter, f (no dry season), w (winter dry), or s (summer dry), and a third symbol (a, b, c, or d [the last subclass exists only for D climates]), indicating the warmth of the summer or the coldness of the winter. The table (Classification of major climatic types according to the Koppen-Geiger-Pohl scheme) gives the specific criteria for the Köppen-Geiger-Pohl system of 1953.

      The Köppen classification has been criticized on many grounds. It has been argued that extreme events, such as a periodic drought or an unusual cold spell, are just as significant in controlling vegetation distributions as the mean conditions upon which Köppen's scheme is based. It also has been pointed out that factors other than those used in the classification, such as sunshine (sunlight) and wind, are important to vegetation. Moreover, it has been contended that natural vegetation can respond only slowly to environmental change, so that the vegetation zones observable today are in part adjusted to past climates. Many critics have drawn attention to the rather poor correspondence between the Köppen zones and the observed vegetation distribution in many areas of the world. In spite of these and other limitations, the Köppen system remains the most popular climatic classification in use today.

      A major contribution to climate grouping was made by the American geographer-climatologist C. Warren Thornthwaite in 1931 and 1948. He first used a vegetation-based approach that made use of the derived concepts of temperature efficiency and precipitation effectiveness as a means of specifying atmospheric effects on vegetation. His second classification retained these concepts in the form of a moisture index and a thermal efficiency index but radically changed the classification criteria and rejected the idea of using vegetation as the climatic integrator, attempting instead to classify “rationally” on the basis of the numerical values of these indices. His 1948 scheme is encountered in many climatology texts, but it has not gained as large a following among a wide audience as the Köppen classification system, perhaps because of its complexity and the large number of climatic regions it defines.

      While vegetation-based climate classifications could be regarded as having relevance to human activity through what they may indicate about agricultural potential and natural environment, they cannot give any sense of how human beings would feel within the various climate types. Terjung's 1966 scheme was an attempt to group climates on the basis of their effects on human comfort. The classification makes use of four physiologically relevant parameters: temperature, relative humidity, wind speed, and solar radiation. The first two are combined in a comfort index to express atmospheric conditions in terms perceived as extremely hot, hot, oppressive, warm, comfortable, cool, keen, cold, very cold, extremely cold, and ultra cold. Temperature, wind speed, and solar radiation are combined in a wind effect index expressing the net effect of wind chill (the cooling power of wind on exposed surfaces) and addition of heat to the human body by solar radiation. These indices are combined for different seasons in different ways to express how humans feel in various geographic areas on a yearly basis. Terjung visualized that his classification would find applicability in medical geography, climatological education, tourism, housing, clothing, and as a general analytical tool.

      Many other specialized empirical classifications have been devised. For example, there are those that differentiate between types of desert and coastal climates, those that account for different rates of rock weathering or soil formation, and those based on the identification of similar agricultural climates.

World distribution of major climatic types
  Classification of major climatic types according to the Koppen-Geiger-Pohl schemeThe following discussion of the climates of the world is based on groupings of Köppen's climatic types. It should be read in conjunction with the table (Classification of major climatic types according to the Koppen-Geiger-Pohl scheme) (for the specific criteria defining each type) and Köppen's climate classification map—> (for the geographic extent of each type).

Type A Climates
      Köppen's A climates are found in a nearly unbroken belt around the Earth at low latitudes, mostly within 15° N and S. Their location within a region in which available net solar radiation is large and relatively constant from month to month ensures both high temperatures (generally in excess of 18 °C [64 °F]) and a virtual absence of thermal seasons (season). Typically, the temperature difference between day and night is greater than that between the warmest and the coolest month, the opposite of the situation in mid-latitudes. The terms winter and summer have little meaning, but in many locations annual rhythm is provided by the occurrence of wet and dry seasons. Type A climates are controlled mainly by the seasonal fluctuations of the trade winds (trade wind), the intertropical convergence zone (ITCZ), and the Asian monsoon (see above Monsoons and Atmospheric pressure and wind: upper-level winds).

Wet equatorial climate (Af)
 Within about 12° latitude of the Equator lies a region of consistently high temperatures (around 30 °C [86 °F]), with plentiful precipitation (150–1,000 cm [59–394 inches]), heavy cloud cover, and high humidity, with very little annual temperature variation. Such regions lie within the influence of the ITCZ in all months; the converging (convergence and divergence), ascending air spawns convectional (convection) thunderstorm activity with much of the rainfall (rain) occurring in late afternoon or early evening when the atmosphere is most susceptible to thunderstorms. While precipitation is profuse in all months, variations do occur in response to the precise location of the ITCZ—drier months result when the ITCZ moves away from the region in question. Other zones of Af climate are found beyond the usual margins of ITCZ activity, in coastal Madagascar, southeast coastal Brazil, and much of Central America and western Colombia, where trade winds blow onshore all year to encounter coastlines backed by mountain barriers that stimulate the formation of precipitation as warm, moist tropical air is forced to ascend and cool. Some of these regions also may receive precipitation from tropical disturbances, including hurricanes (tropical cyclone).

Tropical monsoon and trade-wind littoral climates (Am)
      These climates resemble the Af in most characteristics, with small annual temperature ranges, high temperatures, and plentiful precipitation (often more than Af climates in annual total). They differ from the latter, however, in that they exhibit a short dry season, usually in the low-sun (“winter”) season. The highest temperatures generally occur at the end of this clear spell. Two distinct processes can give rise to Am climate types. The largest areas, mostly in southern and southeastern Asia, result from the Asian monsoon circulation that brings convective and orographic precipitation in the summer when warm, moist, maritime tropical air moves over land to converge into the low-pressure zone north of the Himalayas. In winter, by contrast, cool, dry air diverges out of the Siberian anticyclone to the north, bringing a cooler, drier, and clearer period of variable length. In the Americas and in Africa, Am climates are of the trade-wind variety. These areas receive precipitation on narrow coastal strips through orographic effects as the moist air of the trade winds ascends mountain chains. Seasonal migrations and changes in the intensity of these winds give rise to short, moderately dry seasons. Summer precipitation may be enhanced by tropical disturbances traveling in the trade winds.

Tropical wet-dry climate (Aw)
 This climate has distinct wet and dry seasons, with most of the precipitation occurring in the high-sun (“summer”) season. Total amounts of rainfall are less than in the previous two climate types (50–175 cm [20–69 inches]), most of which occur in convectional thunderstorms. The dry season is longer than in the Am climates and becomes progressively longer as one moves poleward through the region. Temperatures are high throughout the year but show a greater range than Af and Am climates (19–20 °C [66–68 °F] in winter and 24–27 °C [75–81 °F] in summer). Throughout most of the region, the cause of the seasonal cycle is the shift in the tropical circulation throughout the year. During the high-sun season, the ITCZ moves poleward and brings convergent and ascending air to these locations, which stimulates convective rainfall. During the low-sun season, the convergence zone moves off to the winter hemisphere and is replaced by the periphery or core of the subtropical anticyclone (subtropical high), with its subsiding, stable air resulting in a period of dry, clear weather, the intensity and length of which depend on latitude. The subtropical anticyclone occurs in the descending portion of the Hadley cell.

Type B climates
 Arid and semiarid climates cover about a quarter of Earth's land surface, mostly between 50° N and 50° S, but they are mainly found in the 15–30° latitude belt in both hemispheres. They exhibit low precipitation, great variability in precipitation from year to year, low relative humidity, high evaporation rates (when water is available), clear skies, and intense solar radiation.

Tropical and subtropical desert climate (BWh, part of BWk)
 Most of Earth's tropical, true desert (BW) climates occur between 15° and 30° latitude, at the poleward end of the Hadley cell circulation (see atmosphere). These regions are dominated in all months by the subtropical anticyclone, with its descending air, elevated inversions, and clear skies. This is an atmospheric environment that inhibits precipitation. The most extreme arid areas also are far removed from sources of moisture-bearing winds in the interiors of continents and are best developed on the western sides of continents, where the subtropical anticyclone shows its most intense development. An exception to the general tendency for aridity to be associated with subsidence is in the so-called Horn of Africa region (eastern Africa), where the dryness of Somalia is caused more by the orientation of the landmass in relation to the atmospheric circulation. Both the high- and low-sun monsoonal winds blow parallel to the coast, so that moisture-laden maritime air (maritime air mass) can penetrate over land only infrequently. In most low-latitude deserts, cloud cover is uncommon (fewer than 30 days per year have clouds in some areas). Precipitation amounts are mostly in the range 0–25 cm (0–10 inches), although the unreliability of precipitation is more significant than the small totals. Average figures have little meaning; a location with a 10-year mean of 5 cm (2 inches), for example, might have received 50 cm (about 20 inches) in one year as a result of an unusual intrusion of moist air, followed by nine years with no measurable precipitation.

      Temperatures are high, with monthly means in the range 21–32 °C (70–90 °F). Daily temperature variations are extreme. Ranges of 35 °C (63 °F) are not unknown when daytime maxima in excess of 40 °C (104 °F) are followed by a rapid nocturnal temperature drop brought about by the limited capacity of the dry, cloudless desert air to emit infrared radiation to the ground to offset radiation loss from the surface at night. The highest air temperatures recorded on Earth have been in the BWh regions; for example, in shaded, well-ventilated locations, Death Valley in the western United States has reached 57 °C (135 °F), while al-ʿAzīzīyah in Libya has had a recorded high of 58 °C (136 °F). Actual surface temperatures may reach 82 °C (180 °F) on dry sand under intense sunshine.

      An interesting variant of tropical and subtropical deserts are the so-called West Coast Desert areas found on the western coastal margins of the regions discussed above (e.g., in the Sonoran Desert of North America, the Peru and Atacama deserts (Atacama Desert) of South America, and the Sahara [Moroccan part] and Namib deserts (Namib) of Africa). These areas are much cooler than their latitude would suggest (monthly mean temperatures of only 15–21 °C [59–70 °F]), and parts are classified as BWk in Köppen's scheme. The cooling results from airflow off adjacent coastal waters where upwelling of the ocean gives rise to cold currents. Deserts of this sort are subject to frequent fog and low-level clouds; yet they are extremely arid. Some parts of the Atacama Desert, for example, have recorded no precipitation for 20 years.

Tropical and subtropical steppe climate (BSh)
      The low-latitude semiarid (or steppe (Steppe, the)) climate occurs primarily on the periphery of the true deserts treated above. It is transitional to the Aw climate on the equatorward side (showing a summer rainfall maximum associated with the ITCZ and a small annual temperature range) and to the Mediterranean climate on its poleward margin (with a cooler, wetter winter resulting from the higher latitude and mid-latitude frontal cyclone activity). Annual precipitation totals are greater than in BW climates (38–63 cm [15–25 inches]). Yearly variations in amount are not as extreme as in the true deserts but are nevertheless large.

Mid-latitude steppe and desert climate (BSk, part of BWk)
 Although these climates are contiguous with the tropical dry climates of North and South America and of central Asia, they have different origins. Cool true deserts extend to 50° latitude and cool steppes reach nearly 60° N in the Canadian Prairies, well beyond the limits of the subtropical anticyclone. These climates owe their origins to locations deep within continental interiors, far from the windward coasts and sources of moist, maritime air. Remoteness from sources of water vapour is enhanced in some regions (e.g., the Great Plains of the United States) by mountain barriers upwind. Temperature conditions are extremely variable, with annual means decreasing and annual ranges increasing poleward. In the higher latitudes, winters are severely cold, with meager precipitation (much of it in the form of snow) associated with polar and arctic air masses (polar air mass). Summer precipitation is more often convective, arriving in the form of scattered thunderstorm activity brought about by irregular incursions of moist air. Both BWk and BSk climates in the mid-latitudes owe their origins to these mechanisms, but the steppe type tends to be located peripheral to the true desert, either adjacent to the moister C and D climates or at the poleward extent of the range, where reduced evaporation under cooler conditions makes more of the scarce precipitation available as soil moisture for plant growth.

Type C and D climates
 Through a major portion of the middle and high latitudes (mostly from 25° to 70° N and S) lies a group of climates classified within the Köppen scheme as C and D types. Most of these regions lie beneath the upper-level, mid-latitude westerlies (Pacific Ocean) throughout the year, and it is in the seasonal variations in location and intensity of these winds and their associated features that the explanation of their climatic character must be sought. During summer, the polar-front and its jet stream move poleward, and air masses of tropical origin are able to extend to high latitudes. During winter, as the circulation moves equatorward, tropical air retreats and cold polar outbreaks influence weather, even within the subtropical zone. The relative frequency of these air masses of different origins varies gradually from low to high latitude and is largely responsible for the observed temperature change across the belt (which is most marked in winter). The air masses interact in the frontal systems commonly found embedded within the traveling cyclones that lie beneath the polar-front jet stream. Ascent induced by convergence into these low-pressure cells and by uplift at fronts induces precipitation, the main location of which shifts with the seasonal circulation cycle. Other important sources of precipitation are convection, mainly in tropical air, and forced uplift at mountain barriers. Monsoon effects modify this general pattern, while the subtropical anticyclone plays a role in the explanation of climate on the western sides of the continents in the subtropics.

Humid subtropical climate (Cfa, Cwa)
      These climates are found on the eastern sides of the continents between 20° and 35° N and S latitude. Most show a relatively uniform distribution of precipitation throughout the year (the Cfa types), with totals in the range 75–150 cm (30–59 inches). In summer, these regions are largely under the influence of moist, maritime airflow from the western side of the subtropical anticyclonic cells over low-latitude ocean waters. Temperatures are high; the warmest months generally average about 27 °C (81 °F), with mean daily maxima from 30 °C to 38 °C (86 °F to 100 °F) and warm, oppressive nights. Summers are usually somewhat wetter than winters, with much of the rainfall coming from convectional thunderstorm activity; tropical cyclones also enhance warm-season rainfall in some regions. The coldest month is usually quite mild (5–12 °C [41–54 °F]), although frosts (frost) are not uncommon, and winter precipitation is derived primarily from frontal cyclones along the polar front. In North America, the spring and early summer seasons, when the front begins its northward return, are notorious for the outbreak of tornadoes (tornado) associated with frontal thunderstorms along the zone of interaction between tropical and polar air. In eastern and southern Asia, the monsoon influence results in a modified humid subtropical climate (Cwa) that has a clearly defined dry winter when air diverges from the Siberian anticyclone, and the polar front and cyclone paths are deflected around the region. These areas generally lie on the poleward side of Am and Aw climates and exhibit a somewhat larger annual temperature range than Cfa types. Winters are sunny and rather cool. Annual precipitation totals average about 100 cm (39 inches) but vary from 75 to 200 cm (30 to 79 inches).

      Between about 30° and 45° latitude on the western sides of the continents is found a series of climates that show the unusual combination of hot, dry summers and cool, wet winters. Poleward extension and expansion of the subtropical anticyclonic cells over the oceans bring subsiding air to the region in summer, with clear skies and high temperatures. When the anticyclone moves equatorward in winter, it is replaced by traveling, frontal cyclones with their attendant precipitation. Annual temperature ranges are generally smaller than those found in the Cfa climates, since locations on the western sides of continents are not well positioned to receive the coldest polar air, which develops over land rather than over the ocean. Mediterranean climates also tend to be drier than humid subtropical ones, with precipitation totals ranging from 35 to 90 cm (14 to 35 inches); the lowest amounts occur in interior regions adjacent to the semiarid steppe climates. Some coastal locations (e.g., southern California in the western United States) exhibit relatively cool summer conditions and frequent fogs where cold offshore currents prevail. Only in Europe, where the latitude for this climate type fortuitously corresponds to an ocean basin (ocean) (that of the Mediterranean, from which this climate derives its name), does this climate type extend eastward away from the coast for any significant distance.

Marine west coast climate (Cfb, Cfc)
 Poleward of the Mediterranean climate region on the western sides of the continents, between 35° and 60° N and S latitude are regions that exhibit ample precipitation in all months. Unlike their equatorward neighbours, these areas are located beyond the farthest poleward extent of the subtropical anticyclone, and they experience the mid-latitude westerlies and traveling frontal cyclones all year. Precipitation totals vary somewhat throughout the year in response to the changing location and intensity of these storm systems, but annual accumulations generally range from 50 to 250 cm (20 to 98 inches), with local totals exceeding 500 cm (197 inches) where onshore winds encounter mountain ranges. Not only is precipitation plentiful but it is also reliable and frequent. Many areas have rainfall more than 150 days per year, although the precipitation is often of low intensity. Fog is common in autumn and winter, but thunderstorms are infrequent. Strong gales (gale) with high winds may be encountered in winter. These are equable climates with few extremes of temperature. Annual ranges are rather small (10–15 °C or [50–59 °F]), about half those encountered farther to the east in the continental interior at the same latitude. Mean annual temperatures are usually 7–13 °C (45–55 °F) in lowland areas, the winters are mild, and the summers are relatively moderate, rarely having monthly temperatures above 20 °C (68 °F).

      In North and South America, Australia, and New Zealand, north–south mountain ranges backing the west coasts of the landmasses at these latitudes confine the marine west coast climate to relatively narrow coastal strips (but enhance precipitation). By contrast, in Europe the major mountain chains (the Alps and Pyrenees) run east–west, permitting Cfb and Cfc climates to extend inland some 2,000 km (about 1,250 miles) into eastern Germany and Poland.

Humid continental climate (Dfa, Dfb, Dwa, Dwb)
 The D climates are primarily Northern Hemispheric phenomena, since landmasses are absent at the significant latitudes in the Southern Hemisphere. The humid continental subgroup occupies a region between 30° and 60° N in central and eastern North America and Asia in the major zone of conflict between polar and tropical air masses. These regions exhibit large seasonal temperature contrasts with hot summers and cold winters. Precipitation tends to be ample throughout the year in the Df section, being derived both from frontal cyclones and, in summer months, from convectional showers when maritime tropical air pushes northward behind the retreating polar front. Many areas show a distinct summer precipitation maximum because of this convective activity, although more uniform patterns are not uncommon. Severe thunderstorms and tornadoes are an early summer occurrence when the polar front is in the southern margin of the Dfa region. Winter precipitation often occurs in the form of snow, and a continuous snow cover is established for from one to four months in many parts of the region, especially in the north. This snow often arrives in conjunction with high winds from an intense frontal cyclone, giving rise to a blizzard.

      Winters tend to be cold but are subject to occasional frigid or mild spells brought about by periodic incursions of arctic or tropical air. Indeed the changeable nature of weather in all seasons is a characteristic feature of the climate, especially in such areas as the eastern United States and Canada where there are few topographic barriers to limit the exchange of air masses between high and low latitudes. Mean temperatures are typically below freezing from one to several months, and the frost-free season varies from fewer than 150 to 200 days per year. Annual precipitation totals range from 50 to 125 cm (about 20 to 50 inches), with higher amounts in the south of the region and in the uplands.

      In eastern Asia (Manchuria and Korea), a monsoonal variant of the humid continental climate (Dwa, Dwb) occurs. This climate type has a pronounced summer precipitation maximum and a cold, dry winter dominated by continental polar air diverging out of the nearby Siberian anticyclone.

Continental subarctic climate (Dfc, Dfd, Dwc, Dwd)
 North of the humid continental climate, from about 50° to 70° N, in a broad swath extending from Alaska to Newfoundland in North America and from northern Scandinavia to Siberia in Eurasia, lie the continental subarctic climates. These are regions dominated by the winter season, a long, bitterly cold period with short, clear days, relatively little precipitation (mostly in the form of snow), and low humidity. In Asia the Siberian anticyclone, the source of continental polar air, dominates the interior, and mean temperatures 40–50 °C (40–58 °F) below freezing are not unusual. The North American representative of this climate is not as severe but is still profoundly cold. Mean monthly temperatures are below freezing for six to eight months, with an average frost-free period of only 50–90 days per year, and snow remains on the ground for many months. Summers are short and mild, with long days and a prevalence of frontal precipitation associated with maritime tropical air within traveling cyclones. Mean temperatures in summer only rarely exceed 16 °C (61 °F), except in interior regions where values near 25 °C (77 °F) are possible. As a result of these temperature extremes, annual temperature ranges are larger in continental subarctic climates than in any other climate type on Earth, up to 30 °C (54 °F) through much of the area and more than 60 °C (108 °F) in central Siberia, although coastal areas are more moderate. Annual precipitation totals are mostly less than 50 cm (about 20 inches), with a concentration in the summer. Higher totals, however, occur in marine areas near warm ocean currents. Such areas also are generally somewhat more equable and may be designated marine subarctic climates. Areas with a distinct dry season in winter, which results in the Köppen climate types Dwc and Dwd, occur in eastern Siberia, both in the region where the wintertime anticyclone is established and in the peripheral areas subject to dry, divergent airflow from it.

Type E climates
 Köppen's type E climates are controlled by the polar and arctic air masses of high latitudes (60° N and S and higher). These climates are characterized by low temperatures and precipitation and by a surprisingly great diversity of subtypes.

Tundra climate (ET)
 Tundra climates occur between 60° and 75° of latitude, mostly along the Arctic coast of North America and Eurasia and on the coastal margins of Greenland. Mean annual temperatures are below freezing and annual ranges are large (but not as large as in the adjacent continental subarctic climate). Summers are generally mild, with daily maxima from 15 to 18 °C (59 °F to 64 °F), although the mean temperature of the warmest month must be less than 10 °C (50 °F). Days are long (a result of the high latitude), but they are often cloudy. The snow cover of winter melts in the warmer season (though in places with mean annual temperatures of −9 °C [16 °F] or less the ground at depth remains permanently frozen as permafrost); however, frosts and snow are possible in any month. Winters are long and cold (temperatures may be below 0 °C [32 °F] for 6 to 10 months), especially in the region north of the Arctic Circle where, for at least one day in the year, the Sun does not rise. Winter precipitation generally consists of dry snow, with seasonal totals less than in the summer when cyclonic storms that develop along the boundary between the open sea and sea ice yield rainfall. Typical annual totals are less than 35 cm (about 14 inches), but a range from 25 to 100 cm (10 to 39 inches) is possible, with higher totals in upland areas.

Snow and ice climate (EF)
 This climate occurs poleward of 65° N and S latitude over the ice caps of Greenland and Antarctica and over the permanently frozen portion of the Arctic Ocean. Temperatures are below freezing throughout the year, and annual temperature ranges are large but again not as large as in the continental subarctic climates. Winters are frigid, with mean monthly temperatures from −20 °C to −65 °C (–4 °F to –85 °F); the lowest temperatures occur at the end of the long polar night. The EF climate holds the distinction for the lowest recorded temperatures at Earth's surface: the Vostok II research station in Antarctica holds the record for the lowest extreme temperature (−89.2 °C [–129 °F]), while the Plateau Station has the lowest annual mean (−56 °C [–69 °F]). Daily temperature variations are very small, because the presence of snow and ice at the surface refrigerates the air. Precipitation is meager in the cold, stable air (in most cases, 5 to 50 cm [2 to 20 inches]), with the largest amounts occurring on the coastal margins. Most of this precipitation results from the periodic penetration of a cyclone into the region, which brings snow and ice pellets and, with strong winds, blizzards. High winds also occur in the outer portions of the Greenland and Antarctic EF climates, where cold, dense air drains off the higher, central sections of the ice caps as katabatic winds (katabatic wind) (see above Local wind systems (climate)).

Type H climates
 The major highland regions of the world (the Cascades (Cascade Range), Sierra Nevadas (Sierra Nevada), and Rockies (Rocky Mountains) of North America, the Andes (Andes Mountains) of South America, the Himalayas and adjacent ranges and the Tibetan Highlands [or Plateau] of Asia, the eastern highlands of Africa, and the central portions of Borneo and New Guinea) cannot be classified realistically at this scale of consideration, since the effects of altitude and relief give rise to myriad mesoclimates and microclimates. This diversity over short horizontal distances is unmappable at the continental scale. Very little of a universal nature can be written about such mountain areas except to note that, as a rough approximation, they tend to resemble cooler, wetter versions of the climates of nearby lowlands in terms of their annual temperature ranges and seasonality of precipitation. Otherwise, only the most general characteristics may be noted.

      With increasing height, temperature, pressure, atmospheric humidity, and dust content decrease. The reduced amount of air overhead results in high atmospheric transparency and enhanced receipt of solar radiation (especially of ultraviolet wavelength) at elevation. Altitude also tends to increase precipitation, at least for the first 4,000 metres (about 13,100 feet). The orientation of mountain slopes has a major impact on solar radiation receipt and temperature and also governs exposure to wind. Mountains can have other effects on the wind climate; valleys can increase wind speeds by “funneling” regional flows and may generate mesoscale mountain- and valley-wind circulations as well. Cold air also may drain from higher elevations to create “frost pockets” in low-lying valleys. Furthermore, mountains can act as barriers to the movement of air masses, can cause differences in precipitation amounts between windward and leeward slopes (the reduced precipitation on and downwind from lee slopes is called a rain shadow), and, if high enough, can collect permanent snow and ice on their peaks and ridges; the snow line varies in elevation from sea level in the subarctic to about 5,500 metres (about 18,000 feet) at 15–25° N and S latitude.

A. John Arnfield

Climate and life
 The connection (bioclimatology) between climate and life arises from a two-way exchange of mass and energy between the atmosphere and the biosphere. In Earth's early history, before life evolved, only geochemical and geophysical processes determined the composition, structure, and dynamics of the atmosphere. Since life evolved on Earth, biochemical and biophysical processes have played a role in the determination of the composition, structure, and dynamics of the atmosphere. Humans, Homo sapiens, are increasingly shouldering this role by mediating interactions between the biosphere and the atmosphere.

      The living organisms of the biosphere use gases from, and return “waste” gases to, the atmosphere, and the composition of the atmosphere is a product of this gas exchange. It is very likely that, prior to the evolution of life on Earth, 95 percent of the atmosphere was made up of carbon dioxide, and water vapour was the second most abundant gas. Other gases were present in trace amounts. This atmosphere was the product of geochemical and geophysical processes in the interior of Earth and was mediated by volcanic (volcanism) outgassing. It is estimated that the great mass of carbon dioxide in this early atmosphere gave rise to an atmospheric pressure 60 times that of modern times. Today only about 0.035 percent of Earth's atmosphere is carbon dioxide. Much of the carbon dioxide present in Earth's first atmosphere has been removed by photosynthesis, chemosynthesis, and weathering. Currently, most of the carbon dioxide now resides in Earth's limestone sedimentary rocks (sedimentary rock), in coral reefs (coral reef), in fossil fuels (fossil fuel), and in the living components of the present-day biosphere. In this transformation, the atmosphere and the biosphere coevolved through continuous exchanges of mass and energy.

      Biogenic gases are gases critical for, and produced by, living organisms. In the contemporary atmosphere, they include oxygen, nitrogen, water vapour, carbon dioxide, carbon monoxide, methane, ozone, nitrogen dioxide, nitric acid, ammonia and ammonium ions (ion), nitrous oxide, sulfur dioxide (sulfur oxide), hydrogen sulfide, carbonyl sulfide, dimethyl sulfide, and a complex array of non-methane hydrocarbons (hydrocarbon). Of these gases, only nitrogen and oxygen are not “greenhouse gases (greenhouse gas).” Added to this roster of biogenic gases is a much longer list of human-generated gases from industrial, commercial, and cultural activities that reflect the diversity of the human enterprise on Earth.

The Gaia hypothesis
      The notion that the biosphere exerts important controls on the atmosphere and other parts of the Earth system has increasingly gained acceptance among earth and ecosystem scientists. While this concept has its origins in the work of American oceanographer Alfred C. Redfield in the mid-1950s, it was English scientist and inventor James Lovelock that gave it its modern currency in the late 1970s. Lovelock initially proposed that the biospheric transformations of the atmosphere support the biosphere in an adaptive way through a sort of “genetic group selection.” This idea generated extensive criticism and spawned a steady stream of new research that has enriched the debate and advanced both ecology and environmental science. Lovelock called his idea the “Gaia Hypothesis” and defined Gaia as

a complex entity involving Earth's biosphere, atmosphere, oceans, and soil; the totality constituting a feedback of cybernetic systems which seeks an optimal physical and chemical environment for life on this planet.

      The Greek word Gaia, or Gaea, meaning “Mother Earth,” is Lovelock's name for Earth, which is envisioned as a “superorganism” engaged in planetary biogeophysiology. The goal of this superorganism is to produce a homeostatic (homeostasis), or balanced, Earth system. The scientific process of research and debate will eventually resolve the issue of the reality of the “Gaian homeostatic superorganism,” and Lovelock has since revised his hypothesis to exclude goal-driven genetic group selection. Nevertheless, it is now an operative norm in contemporary science that the biosphere and the atmosphere interact in such a way that an understanding of one requires an understanding of the other. Furthermore, the reality of two-way interactions between climate and life is well recognized.

The evolution of life and the atmosphere
      Life on Earth began at least as early as 3.5 billion years ago during the middle of the Archean Eon (about 4 billion to 2.5 billion years ago). It was during this interval that life first began to exercise certain controls on the atmosphere. The atmosphere's prebiological state is often characterized as being rich in water vapour and carbon dioxide. Though some nitrogen was also present, little if any oxygen was available. Chemical reactions with hydrogen sulfide, hydrogen, and reduced compounds of nitrogen and sulfur precluded any but the shortest lifetime for free oxygen in the atmosphere. As a result, life evolved in an atmosphere that was reducing (reduction) (high hydrogen content) rather than oxidizing (oxidation–reduction reaction) (high oxygen content). In addition to their chemically reducing character, the predominant gases of this prebiotic atmosphere, with the exception of nitrogen, were largely transparent to incoming sunlight but opaque to outgoing terrestrial infrared radiation. As a result, these gases are called, perhaps improperly, greenhouse gases (greenhouse effect) (see greenhouse effect) because they are able to slow the release of outgoing radiation back into space.

      In the Archean Eon, the Sun produced as much as 25 percent less light than it does today; however, Earth's temperature was much like that of today. This is possible because the greenhouse gas-rich Archean atmosphere was effective in retarding the loss of terrestrial radiation to space. The resulting long residence time of energy within the Earth-atmosphere system resulted in a warmer atmosphere than would have been possible otherwise. The average temperature of Earth's surface in the early Archean Eon was warmer than the modern global average. It was, according to some sources, probably similar to temperatures found in today's tropics. Depending on the amount of nitrogen present during the Archean Eon, it has been suggested that the atmosphere may have held more than 1,000 times as much carbon dioxide than it does today.

      Archean organisms included photosynthetic and chemosynthetic bacteria, methane-producing bacteria, and a more primitive group of organisms now called the “ archaea” (a group of prokaryotes (prokaryote) more related to eukaryotes (eukaryote) than to bacteria and found in extreme environments). Through their metabolic (metabolism) processes, organisms of the Archean Eon slowly changed the atmosphere. Hydrogen rose from trace amounts to about 1 part per million (ppm) of dry air. Methane concentrations increased from near zero to about 100 ppm. oxygen increased from near zero to 1 ppm, whereas nitrogen concentrations rose to encompass 99 percent of all atmospheric molecules excluding water vapour. Carbon dioxide concentrations decreased to only 0.3 percent of the total; however, this was nearly 10 times the current concentration. The composition of the atmosphere, its radiation budget, its thermodynamics, and its fluid dynamics were transformed by life from the Archean Eon.

      American geochemist Robert Garrels calculated that, in the absence of life and given the burial rate of carbon in rocks (rock), oxygen would be unavailable to form water, and free hydrogen would be lost to space. Without the presence of life and compounded by this loss of hydrogen, there would be no oceans (ocean), and Earth would have become merely a dusty planet by the middle of the Archean Eon. By the end of the Archean Eon 2.5 billion years ago, both the pigment chlorophyll and photosynthetic organisms had evolved such that the production of oxygen increased rapidly. The atmosphere became transformed from a reducing atmosphere with carbon dioxide, limited oxygen, and anaerobic organisms (that is, life-forms that do not require oxygen for respiration) in control to one with an oxidizing atmosphere that was rich in oxygen, poor in carbon dioxide, and dominated by aerobic organisms (that is, life-forms requiring oxygen for respiration).

      With the decline in carbon dioxide and a rise in oxygen, the greenhouse warming capacity of Earth's atmosphere was sharply reduced; however, this happened over a period of time when the energy produced by the Sun increased systematically. These compensating changes resulted in a relatively constant planetary temperature over much of Earth's history.

The role of the biosphere in the Earth-atmosphere system
The biosphere and Earth's energy budget
 Biogenic gases in the atmosphere play a role in the dynamics of Earth's planetary radiation budget, the thermodynamics of the planet's moist atmosphere, and, indirectly, the mechanics of the fluid flows that are Earth's planetary wind systems. In addition, human cultural and economic activities add a new dimension to the relationship between the biosphere and the atmosphere. While humans are biologically trivial compared with bacteria in the exchange of gases with the atmosphere, chemical compounds produced from human industrial activities and other economic enterprises are changing the gaseous composition of the atmosphere in climatically significant ways. The largest changes involve the harvesting of ancient carbon stores. This organic material has been transformed into fossil fuels (fossil fuel) ( coal, petroleum, natural gas, and others) by geologic processes acting upon the remains of plants (plant) and animals (animal) over many millions of years. Different forms of carbon may be burned and thus used as energy sources. In so doing, organic carbon is converted into carbon dioxide. Additionally, humans are also burning trees, grasses, and other biomass for cooking purposes and clearing the land for agriculture and other activities. The combination of burning both fossil fuels and biomass is enriching the atmosphere with carbon dioxide and adding to the essential reservoir of greenhouse gases (see global warming).

      Earth's atmosphere is largely transparent to sunlight. Of the sunlight absorbed by the entire Earth-atmosphere system, about one-third is absorbed by the atmosphere and two-thirds by Earth's surface. Sunlight is absorbed by the molecules of the atmosphere, by cloud droplets, and by dust and debris. Though oxygen and nitrogen make up nearly 99 percent of the atmosphere, these diatomic molecules do not vibrate in a way that permits them to absorb terrestrial radiation. They are largely transparent to outgoing terrestrial radiation as well as to incoming solar radiation.

      Over the continents, the surface cover of vegetation is the principal absorbing medium of Earth's surface, although other surfaces such as bare rock, sand, and water also absorb solar radiation. At night, absorption at the surface (that is, below 1.2 metres [4 feet]) is reradiated, in the form of long-wave infrared radiation, away from Earth's surface back toward space. Most of this infrared radiation is absorbed by the principal biogenic trace gases of the atmosphere—the so-called greenhouse gases (greenhouse gas): water vapour, carbon dioxide, and methane. Without these biogenic greenhouse gases, Earth would be 33 °C (59 °F) colder on average than it is. A moderate-emission scenario from the 2007 Intergovernmental Panel on Climate Change (IPCC) report predicts that the continued addition of greenhouse gases from fossil fuels will increase the average global temperature by between 2.3 and 4.3 °C (4.1 and 7.7 °F) over the next century. Other scenarios, predicting greater greenhouse gas emissions, forecast even greater global warmin (global warming)g.

The cycling of biogenic atmospheric gases
       Average Composition of the Atmosphere Average Composition of the AtmosphereThe cycling of oxygen, nitrogen, water vapour, and carbon dioxide, as well as the trace gases—methane, ammonia, various oxides of nitrogen and sulfur, and non-methane hydrocarbons—between the atmosphere and the biosphere results in relatively constant proportions of these compounds in the atmosphere over time. Without the continuous generation of these gases by the biosphere, they would quickly disappear from the atmosphere.

 The carbon cycle, as it relates to the biosphere, is simple in its essence. Inorganic carbon (carbon dioxide) is converted to organic carbon (the molecules of life). To complete the cycle, organic carbon is then converted back to inorganic carbon. Ultimately, the carbon cycle is powered by sunlight as green plants and cyanobacteria (blue-green algae) ( blue-green algae) use sunlight to split water into oxygen and hydrogen and to fix carbon dioxide into organic carbon. Carbon dioxide is removed from the atmosphere, and oxygen is added. Animals engage in aerobic respiration, in which oxygen is consumed and organic carbon is oxidized to manufacture inorganic carbon dioxide. It should be noted that chemosynthetic bacteria, which are found in deep-ocean and cave ecosystems (ecosystem), also fix carbon dioxide and produce organic carbon. Instead of using sunlight as an energy source, these bacteria rely on the oxidation of either ammonia or sulfur.

 The carbon cycle is fully coupled to the oxygen cycle. Each year, photosynthesis fixes carbon dioxide and releases 100,000 megatons of oxygen to the atmosphere. Respiration by animals and living organisms consumes about the same amount of oxygen and produces carbon dioxide in return. Oxygen and carbon dioxide are thus coupled in two linked cycles. On a seasonal basis, an enrichment of atmospheric carbon dioxide occurs in the winter half of the year, whereas a drawdown of atmospheric carbon dioxide takes place during the summer.

 The nitrogen cycle begins with the fixing of inorganic atmospheric nitrogen (N2) into organic compounds (organic compound). These nitrogen-containing compounds are used by organisms and, through the process of denitrification, are converted back to inorganic atmospheric nitrogen. Ammonia and ammonium ions are the products of nitrogen fixation and may be incorporated into some organisms as organic nitrogen-containing molecules. In addition, ammonium ions may be oxidized to form nitrites (nitrite), which can be further oxidized by nitrifying bacteria into nitrates (nitrate). Though both nitrites and nitrates may be used to make organic nitrogen-containing molecules, nitrates are especially useful for plant growth and are key compounds that support both terrestrial and aquatic food chains (food chain). Nitrates may also be denitrified by bacteria to produce nitrogen gas. This process completes the nitrogen cycle.

      The nitrogen cycle is coupled to both the carbon and oxygen cycles. Seventy-eight percent of the gases of the atmosphere by volume is diatomic nitrogen. Diatomic nitrogen is the most stable of the nitrogen-containing gases of the atmosphere. Only 300 megatons of nitrogen must be produced each year by denitrifying bacteria to account for losses. These losses mostly occur during lightning discharges and during nitrogen-fixation activities by blue-green algae and nitrogen-fixing bacteria. (The latter are found in the root nodules of certain plants called legumes (legume).) Nitrogen-fixing bacteria use atmospheric nitrogen to produce oxides of nitrogen. Denitrifying bacteria convert nitrates in soils and wetlands (wetland) to nitrogen gas, which is then returned to the atmosphere.

      Nitrous oxide occurs in trace amounts (0.3 ppm) in the atmosphere. Between 100 and 300 megatons of nitrous oxide are produced by soil and marine bacteria each year to maintain this concentration. In the atmosphere, nitrous oxide is short-lived because it is quickly broken down by ultraviolet light (ultraviolet radiation). nitric oxide (NO), a minor contributor in the breakdown of stratospheric ozone, is the by-product of this reaction.

      Like nitrous oxide, ammonia is also produced in soils and marine waters by bacteria and escapes into the atmosphere. Nearly 1,000 megatons are added to the atmosphere each year. Ammonia decreases the acidity of precipitation and serves as a nutrient for plants when returned to the land via precipitation.

      There are six major sulfur-containing biogenic atmospheric gases that are part of the sulfur cycle. They include hydrogen sulfide, carbon disulfide (CS2), carbonyl sulfide (COS), dimethyl sulfide (DMS; C2H6S), dimethyl disulfide ([CH3S]2), and methyl mercaptan (CH3SH). Sulfur dioxide (SO2) is an oxidation product of these sulfur gases, and it is also added to the atmosphere by volcanoes (volcano), burning biomass, and anthropogenic sources (i.e., smelting metals and coal ignition). SO2 is removed from the atmosphere and returned to the biosphere in rainfall (rain). The increased acidity of rain and snow from anthropogenic additions of SO2 and oxides of nitrogen is often referred to as “acid precipitation (acid rain).” The acidity of this precipitation and other phenomena, such as “acid fog,” is partly cancelled by the release of ammonia in the atmosphere.

      The concentration of methane at any one time in the atmosphere is only about 1.7 ppm. Though only a trace gas, it is highly reactive and plays a key role in the chemical reactions that control the composition of the atmosphere. Methanogenic bacteria in wetland sediments decompose organic matter and release 1,000 megatons of gaseous methane to the atmosphere per year. In the lower atmosphere, methane reacts with oxygen to produce water and carbon dioxide. Each year 2,000 megatons of oxygen are removed from the atmosphere by this mechanism. This loss of oxygen must be replaced by photosynthesis. Some methane reaches the upper stratosphere, where its interaction with oxygen is a major source of upper stratospheric moisture. Within wetlands, bacteria produce methyl halide compounds ( methyl chloride [CH3Cl] and methyl iodide [CH3I] gases), whereas these same methyl halides are produced in forests by fungi (fungus). These gases, upon reaching the stratosphere, regulate the production of stratospheric ozone by contributing to its natural breakdown (see ozonosphere). Without the continual production of methane by methanogenic bacteria, the oxygen concentration of the atmosphere would increase by 1 percent in only 12,000 years. Dangerously high levels of oxygen in the atmosphere would greatly increase the incidence of wildfires (wildland fire). If the oxygen concentration of Earth's atmosphere rose from its current concentration of 21 percent to 25 percent, even damp twigs and grass would easily ignite. Non-methane hydrocarbons of terrestrial origin are generally well mixed in the free atmosphere above the planetary boundary layer (PBL; see below). These organic particles weaken incoming solar radiation as it passes through the atmosphere, and reductions of 1 percent have been recorded.

biosphere controls on the structure of the atmosphere
      Because the biosphere plays a key role in the flux of energy from the surface to the atmosphere, it also contributes to the structure of the atmosphere. Three major fluxes are important: the direct transfer of heat from the surface to the atmosphere by conduction (thermal conduction) and convection (sensible heating), the energy flux to the atmosphere carried by water vapour via evaporation (vaporization) and transpiration from the surface ( latent heat energy), and the flux of radiant energy from the surface to the atmosphere (infrared terrestrial radiation). These fluxes differ in the altitude at which the heating of the air takes place and thus contribute to the thermal structuring of the atmosphere. Sensible heating primarily warms the planetary boundary layer (PBL) of the atmosphere. In marine areas, the PBL occurs in the lowest 1 km (3,300 feet); in heavily vegetated areas, the PBL occurs in the lowest 1 to 2 km (3,300 to 6,600 feet); and in arid regions, it occurs in the lowest 4 or 5 km (13,100 to 16,400 feet) of the atmosphere. In contrast, the latent heat of the atmosphere is released when the water vapour is converted into cloud droplets by condensation. Heating by latent energy release generally occurs above the PBL.

      On the other hand, heating of the atmosphere by radiation from the surface depends on the density of the atmosphere and its water vapour content. Radiative heating from the surface declines with increasing altitude. The availability of water to evaporate from the surface limits the sensible heating of the air near the surface and so limits the maximum daytime surface air temperature (see below).

Biosphere controls on the planetary boundary layer
 The top of the planetary boundary layer (PBL) can be visually marked by the elevation of the base of the clouds. In addition, the PBL can also be denoted by a thin layer of haze often seen by passengers aboard airplanes during takeoff from airports. During the day, the air within the PBL is thoroughly mixed by convection induced by the heating of Earth's surface. The thickness of the PBL depends on the intensity of this surface heating and the amount of water evaporated into the air from the biosphere. In general, the greater the heating of the surface, the deeper the PBL. Over deserts, the PBL may extend up to 4 or 5 km (13,100 or 16,400 feet) in altitude. In contrast, the PBL is less than 1 km (0.6 mile) thick over ocean areas, since little surface heating takes place there because of the vertical mixing of water. The wetter the air advected into the region and the greater the additional water added by evaporation and transpiration, the lower the height of the top of the PBL. For every 1 °C (1.8 °F) increase in daily maximum surface temperature for a well-mixed PBL, the top of the PBL is elevated 100 metres (about 325 feet). In New England forests during the days following the spring leafing, it has been shown that the top of the PBL is lowered to between 200 and 400 metres (650 and 1,300 feet). By contrast, during the months before the leafing out, the PBL thickens from solar heating as the sun rises higher in the sky and day length increases.

      If convective mixing of the air in the PBL is vigorous, convection currents may penetrate through the temperature inversion at the top of the PBL. The cooling of the lifting air initiates the condensation of water vapour and the development of miniscule particles of liquid water called cloud droplets. The small clouds just above the PBL are known as planetary boundary layer clouds. These clouds scatter direct sunlight. As the ratio of diffuse sunlight to direct beam sunlight increases, greater levels of photosynthetic productivity are favoured in the biosphere below. The result is a dynamic synergy between the atmosphere and biosphere.

      The landscapes of most human-dominated ecosystems are decidedly “patchy” in their geography (see below). Cities, suburbs, fields, forests, lakes, and shopping centres both heat and evaporate water into the air of the PBL according to the nature of the surfaces involved. Convection and the prospect of breaking through the top of the PBL vary markedly across such heterogeneous landscapes. These upward and downward currents or vertical eddies (eddy) within the PBL transfer mass and energy upward from the surface. The frequency, timing, and strength of convective weather elements, including thunderstorms (thunderstorm), vary according to the patchiness of the land use and land cover pattern of the area. In general, the greater the patchiness of the landscape and the earlier the hour in the day, the more frequent and more intense these rain-producing systems become.

      In the absence of an organized storm in the region, the air above the PBL sinks gently and the air below lifts. At the top of the PBL, a small inversion, where temperatures increase with height, develops. This inversion essentially becomes a stable layer in the atmosphere. Emissions from the biosphere below are thus contained within the PBL and may build up below this layer over time. Consequently, the PBL may become quite turbid, hazy, or filled with smog.

 When the sinking from above is vigorous, the PBL inversion grows in thickness. This situation has the effect of hindering the development of thunderstorms, which depend on rapidly rising air. This often occurs over southern California, and thus the chance of thunderstorms forming there is small. Emissions from both the biosphere and from anthropogenic activities accumulate in this part of the atmosphere, and pollution may build up to such an extent that health warnings may be required. In locations free of temperature inversions, convection processes are strong enough, particularly during the summer months, that emissions are scavenged and quickly lifted by thunderstorms to regions high above the PBL. Often, acidic compounds from these emissions are returned to the surface in the precipitation that falls.

Biosphere controls on maximum temperatures (temperature) by evaporation and transpiration
      Solar radiation is converted to sensible and latent heat at Earth's surface. A change in sensible heat results in a change in the temperature of a medium, whereas energy stored as latent heat is used to drive a process, such as a phase change in a substance from its liquid to its gaseous state, and does not produce a change in temperature. Thus, the daily maximum surface temperature at a given location is dependent on the amount of radiant energy converted to sensible heat. water available for evaporation increases latent heating by adding water vapour to the atmosphere. As a result, relatively little energy remains to heat the air, and thus the sensible heating of the air near the ground is minimized. In addition, daily maximum temperatures are not as high in locations with strong latent heating.

      As day length increases from winter to summer, sensible heating and maximum surface temperatures rise. In the U.S. Midwest, prior to the leafing out of vegetation (leaf) in the springtime and the resulting rise in evaporation and transpiration, sensible heating causes an average increase in maximum surface temperatures of only about 0.3 °C (0.5 °F) per day. The process of leaf production creates a surge in evaporation and transpiration and results in increased latent heating and reduced sensible heating. After leafing, since most of the available thermal energy is used to convert liquid water to water vapour rather than to heat the air, the average day-to-day rise in daily maximum temperatures is reduced to about 0.1 °C (0.2 °F) per day.

      This effect extends upward through the atmosphere. Prior to leafing out, the one-kilometre-thick layer occurring between the 850-to-750-millibar pressure level (which typically occurs between 1,650 and 2,750 metres [5,400 and 9,000 feet]) in the Midwest warmed at the rate of 0.1 °C (0.2 °F) per day. Following leafing out, the warming rate fell to 0.02 °C (0.04 °F) per day. Scientists have used computer models of the atmosphere to study the effect of transpiration from vegetation on maximum surface air temperatures. In these models, the variable controlling transpiration by vegetation was “turned off,” and the character of the resulting modeled climate was studied. By subtracting the effect of transpiration, temperatures in central North America and on the other continents were predicted to equilibrate at a very hot 45 °C (113 °F). Such warming is nearly realized in desert areas where moisture is unavailable for transpiration.

Biosphere controls on minimum temperatures (temperature)
      During the late 1860s, British experimental physicist John Tyndall (Tyndall, John), based on his studies of the infrared radiation absorption by atmospheric gases, concluded that nighttime minimum temperatures were dependent on the concentration of trace gases in the atmosphere. Of these gases, water vapour had the greatest impact. To emphasize the significance of water vapour on decreases in air temperature during the night, he wrote that if all the water vapour in the air over England was removed even for a single night, it would be “attended by the destruction of every plant which a freezing temperature could kill.” As a result, it follows that the greater the water content of the atmosphere, the lower the radiative loss of energy to the sky and the less the surface atmosphere is cooled. Thus, locations with substantial amounts of water vapour experience reduced nocturnal cooling.

 Water vapour in the atmosphere also limits the extent to which temperatures fall at night. This limiting temperature is known as the dew point, which is defined as the temperature at which condensation begins. Over North America east of the 100th meridian (a line of longitude traditionally dividing the moist eastern part of North America from drier western areas), average nighttime minimum temperatures are within a degree or two of the dew point temperature. Upon nocturnal cooling, the dew point is reached, condensation begins, and latent energy is converted to heat. Additional temperature falls are retarded by this release of heat to the atmosphere. A significant fraction of the water in the atmosphere over the continents comes from the evaporation of water from soils and the transpiration from vegetation. Transpired water directly moderates temperature by increasing humidity and thus raising the dew point. As a consequence, the amount of outgoing terrestrial radiation released to space is reduced. This results in the elevation of the minimum temperature of the air above what it would otherwise be.

 The effect of spring leafing (leaf) on the buildup of humidity in the lower atmosphere has received the attention of researchers in recent years. In the late 1980s, American climatologists M.D. Schwartz and T.R. Karl used the superimposed epoch method to study the climate before and after the leafing out of lilac plants in the spring in the U.S. Midwest. (This method uses time series data from multiple locations, which can be compared to one another by adjusting each data set around the respective onset date of lilac blooming.) In the illustration, the x-axis marks days before and after leafing, whereas the y-axis shows the related changes in the vapour pressure of the atmosphere. A second y-axis follows the day-by-day changes in minimum temperatures. Prior to the average date of leafing, the atmospheric humidity (vapour pressure) is relatively constant and minimum temperatures hover near freezing. At leafing, there is an abrupt increase in atmospheric humidity. Following leafing, daily minimum temperatures also increase abruptly. Although frosts are possible until June 10th in many parts of the Midwest, the chances of frost decline as the atmosphere is humidified.

Climate and changes in the albedo of the surface
      The amount of solar energy available at the surface for sensible and latent heating of the atmosphere depends on the albedo, or the reflectivity, of the surface. Surface albedos vary by location, season, and land cover type. The albedo of unvegetated ground devoid of snow ranges from 0.1 to 0.6 (10 to 60 percent), while the albedo of fully forested lands ranges from 0.08 to 0.15. An increase of 0.1 in regional albedo has been associated with a 20 percent decline in rainfall events connected with thunderstorms. Equivalent reductions in both evaporation and transpiration have also been reported in areas with sudden increases in albedo.

  The greatest changes in albedo occur in regions undergoing desertification and deforestation. Depending on the albedo of the underlying soil, reductions in vegetative land cover may give rise to albedo increases of as much as 0.2. Model studies of the vegetative zone known as the Sahel in Africa reveal that albedo increased from 0.14 to 0.35 due to desertification occurring during the 20th century. This coincided with a 40 percent decrease in rainfall. In addition, it is likely that the clearing of forests and prairies for agricultural crops over the past several hundred years has altered the albedo of extensive regions of the middle latitudes.

      Contemporary agricultural practices give rise to large variations in albedo from season to season as the land passes through the cycle of tilling, planting, crop growth, and harvest. At larger scales, an agricultural mosaic often emerges as each different plot of ground is covered by plantings of a single species. Viewed from the air, landscapes in the middle latitudes appear as a heterogeneous mix of forests, grasslands, meadows, water bodies, farmlands, wetlands, and urban types. The resultant patchiness in the landscape produces a patchiness in surface albedo. The mosaic of land use types creates a mix in the fluxes of sensible and latent heat to the atmosphere. Such changes to the heat flux have been shown to cause changes in the timing, intensity, and frequency of summer thunderstorms.

The effect of vegetation patchiness on mesoscale climates
      The establishment of vegetation bands or patches 50 to 100 km (30 to 60 miles) in width in semiarid regions could increase atmospheric convection and precipitation beyond that expected over areas of uniform vegetation. This convection creates spatial differences in the upward and downward wind velocities and contributes to the development of mesoscale (20 to 200 km [12 to 120 miles]) circulation in the atmosphere (see Upper-level winds: Characteristics (climate)). For example, when creating models for forecasting atmospheric conditions on the Great Plains and along the Front Range of the Rocky Mountains, the mix of land cover and vegetation types must be specified to properly relate the fluxes of momentum and sensible and latent heat to the larger-scale circulation of the atmosphere. Proper calculations are also necessary to estimate rainfall. In addition, the specific location and hour of the day that thunderstorms occur depend on the heterogeneity of the vegetation cover of this region. Field observations have shown that the heterogeneity of surface roughness (small-scale irregularities in topography), soil moisture, forest coverage, and transpiration affect the location and pace of the formation of convective clouds and rainfall. Both convection and thunderstorm development tend to occur earlier in the day in heterogeneous landscapes.

Biosphere controls on surface friction and localized winds
      Averaged annually over Earth's entire surface, the Sun provides about 345 watts per square metre of energy. About 30 percent of this energy is reflected away to space and is never used in the Earth-atmosphere system. Of that which remains, a little less than 1 percent (3.1 watts per square metre) accelerates the air by generating winds (wind). An equal amount of energy must eventually be lost, or else wind speeds would perpetually increase.

      Earth as a thermodynamic system is dissipative—the mechanical energy of the winds is eventually converted to heat through friction. Over the continents, it is the combination of terrain and the veneer of vegetation that offers the frictional roughness to dissipate the surface winds and convert this kinetic energy into heat. Marine winds approaching the British Isles average about 12 metres per second (27 miles per hour), but they are decelerated to 6 metres per second (13 miles per hour) because of the friction of the landscape's surface shortly after the winds make landfall. Without vegetation cover, the continents would offer much less friction to the wind, and wind speeds in unvegetated landscapes would be nearly twice as fast as those in vegetated landscapes.

      The correct specification of Earth's surface roughness due to vegetation, for use in computer models of the atmosphere, is critical to proper model performance. If the height of the terrain and vegetation are not specified correctly, the patterns of Earth's winds, global geography, and rainfall will be poorly modeled. When modeling newly desertified areas, such as the Sahel, it is important to understand that desertification creates vegetation of lower stature and thus lower surface roughness values. As a result, both wind velocities and wind direction could change from previous patterns over landscapes with taller vegetation.

      The extent of this impact of the biosphere on the atmosphere is revealed in climate model studies. One such study modeled the influence of reduced vegetation on surface roughness over the Indian subcontinent and provided evidence for a weaker monsoon and reduced rainfall. Given that much of the northwest third of India underwent a severe desertification and cultural collapse near the beginning of historical times, the role cultures play in vegetation reduction and climate change should not be ignored.

      The vegetation cover of the continents is not passive in response to the winds. Greenhouse-grown trees subjected to mechanical forces designed to mimic the winds lay down new woody tissue called “reaction wood,” which results in a stiffer tree over time. This material helps trees become more resilient and offer more frictional resistance to wind. This negative feedback, where increased winds result in stiffer vegetation and thereby subsequently reduced wind speeds, might well apply at the global scale by balancing the energy used to heat and accelerate the air (3.1 watts per square metre) with the surface friction needed to dissipate it.

Biosphere impacts on precipitation processes

Cloud condensation nuclei
      The formation and subsequent freezing of cloud droplets depend on the presence of cloud condensation nuclei (condensation nucleus) and ice nuclei, respectively. Significantly, the biosphere is a major source of both of these kinds of nuclei. Over the continents, condensation nuclei are readily available and are of biogenic as well as anthropogenic origin. Examples of condensation nuclei include sea salt, small soil particles, and dust.

      As atmospheric convection increases with the heating of the day, cloud condensation nuclei are mixed into and above the planetary boundary layer and into the troposphere. In the bottom 0.5 km (the lowest 1,600 feet or so) of the atmosphere, nuclei typically number 2.2 × 1010 per cubic metre. In the next 0.5 km (between 1,600 and 3,300 feet) above, half as many nuclei are found. The number of condensation nuclei continues to decline with increased altitude. Furthermore, in general, the number of nuclei in the air over land is 10 times higher than over the oceans.

      Cloud condensation nuclei are generally abundant. They do not limit cloud formation over the continents; however, low numbers of condensation nuclei over the oceans may limit cloud formation there. In addition to natural sources, particulates from fuel combustion and sulfur dioxide gas resulting from high sulfur fuels also contribute to the load of condensation nuclei over the continents. Both the number and kind of condensation nuclei present in the atmosphere affect the cloudiness and the brightness of clouds in a given region. In this way, condensation nuclei play a significant role in determining both regional and global albedo.

      There is a type of condensation nuclei that forms in the marine air over the margins of continents. Though these nuclei are often few in number, they play a large role in cloud formation near the coasts of continents and may contribute significantly to both planetary albedo and global average temperature. Typically, sources of condensation nuclei in marine air are sulfate aerosols formed from the biogenic production of dimethyl sulfide (DMS) by marine algae. Given that DMS production increases with sea surface temperatures, a negative feedback may result. The central idea in this feedback hypothesis is that warmer waters result in the increased production of condensation nuclei by phytoplankton and thus produce more clouds. Increased cloudiness shades the ocean surface and results in lower temperatures that limit condensation nuclei production. It is estimated that a 30 percent increase in marine condensation nuclei would increase planetary albedo by 0.005 (0.5 percent) or produce a 0.7 percent reduction in solar radiation and a planetary average temperature decrease of 1.3 °C (2.3 °F). The sensitivity of this negative feedback on planetary temperatures remains in active debate.

Biogenic ice nuclei
      As water vapour condenses onto condensation nuclei, the droplets grow in size. Growth proceeds at relative humidity as low as 70 percent, but the rate of growth is very slow. Growth by condensation is most rapid where the air is slightly supersaturated with water vapour. At this point, cloud droplets typical of the size of fog droplets arise. Should temperatures fall to the level where freezing begins, the temperature difference between the droplet and the surrounding air (the vapour pressure deficit) strongly favours rapid condensation into the crystalline lattice of an ice particle. Ice particles that grow rapidly soon reach sizes where they begin to fall. As they fall, they collide and merge with smaller droplets and thereby grow larger.

      The formation of ice is of critical importance. A droplet of pure water, such as distilled water, will automatically freeze in the atmosphere at a temperature of –40 °C (–40 °F). Freezing at warmer temperatures requires a substance upon which ice crystallization can take place. The common clay mineral kaolinite, a contaminant of the droplet, raises this freezing point to around –25 °C (–13 °F). Furthermore, silver iodide, often used in cloud seeding to encourage rainfall, and sea salts also cause ice to form at –25 °C. Freezing at still warmer temperatures is most common with biogenic ice nuclei. Upon ice formation, heat energy on the order of 80 calories per gram of water frozen are released. This energy increases the sensible heat of the air and causes the air to become more buoyant. The process of ice formation encourages convection, cloudiness, and precipitation from clouds.

      The decomposition of organic matter is a major source of biogenic ice nuclei. Ice crystal formation has been shown to occur at temperatures as warm as –2 to –3 °C (28.5 to 26.6 °F) when biogenic ice nuclei are involved. The common freezing temperature for biogenic nuclei varies systematically according to biome and latitude. The coldest freezing-temperature nuclei occur above the tropics, whereas the warmest occur above the Arctic. Freezing produces greater buoyancy of the particles and helps them to reach higher vertical velocities within the clouds. The vertical motions and the larger droplet size that occur with biogenic materials favour the charge separation needed to produce lightning. Subsequently, oceanic areas with few biogenic ice nuclei are also areas of low lightning frequency. The production of biogenic nuclei from organic matter decomposition is greatest during the warm months when bacterial decomposition is greatest.

Recycled rainfall
      The water that is transpired into the atmosphere from the biosphere is eventually returned to the surface as precipitation. This vegetation-transpiration component of the hydrologic cycle is referred to as “recycled rainfall.” While the oceans are the major source of atmospheric water vapour and rainfall, water from plant transpiration is also significant. For example, in the 1970s and '80s, analyses performed by American meteorologist Michael Garstang on the city of Manaus, Braz., in the Amazon (Amazon Rainforest) basin revealed that around 20 percent of the precipitation came from water transpired by vegetation; the remaining 80 percent of this precipitation (an estimate made by German American meteorologist Heinz Lettau in the 1970s) was generated by the Atlantic Ocean. Isotopic studies of rainwater collected at various points in the Amazon basin indicated that nearly half of the total rain came from water originating in the ocean and half transpired through the vegetation. Evidence of the proportion of transpired water in rainfall reaching as high as 88 percent has been reported for the Amazon foothills of the Andes. General climate circulation models indicate that, without transpired water from plants, rainfall in the central regions of the continents would be greatly reduced. As a general rule, the farther the distance from oceanic water sources, the higher the fraction of rainwater originating from transpiration.

Climate, humans, and human affairs
 In the late 1960s and early '70s, climatologists envisioned the start of a new ice age because it was becoming clear that a cycle of planetary cooling was detected in the post-1940s record. The Central Intelligence Agency of the United States commissioned studies of the global political stresses that would ensue with a 1 °C (1.8 °F) temperature decline. The central question revolved around the locations of the political hot spots in a cooling world. The results of these studies were published under the title The Weather Conspiracy. In the late 1970s, Harvard University's Center for International Affairs addressed these issues in a book by English diplomat and environmentalist Crispin Tickell titled Climatic Change and World Affairs. Tickell sounded a warning:A shift of 2 °C in mean temperatures leads either to ice ages or to melting of the polar ice caps, either of which would destroy much of present civilization.

      In the late 1970s the global warming concerns arising from the burning of fossil fuels were still a decade away, but the Harvard report provided the impetus to research the possible links between the burning of fossil fuels and global warming. In academia during this time, climatologists and historians began working together to reexamine the past connections between climate and history.

      Since prehistoric times, humans have altered the land cover of the continents to suit their economic and cultural enterprises. In so doing, they have recast the two-way balance of mass and energy exchange between the atmosphere and the biosphere. Since the characteristics of Earth's climate in the absence of cities, agriculture, and other human land uses are unknown, comparing the dynamics of the present Earth-atmosphere system with that of preurban and preagricultural times is very difficult.

      In any case, humans are subject to the same climatic variations (climate change) and changes that affect other life-forms. Food and fibre resources are climatically sensitive and vary in price accordingly. The cost of hauling coal by railroad depends on the temperature of the rails and on the weight gain that occurs when coal is wetted by rainfall. The consumption of fossil fuels for heating and cooling purposes is also climate-dependent. Damage from hurricanes (tropical cyclone), floods (flood), droughts (drought), snowstorms (blizzard), tornadoes (tornado), and other weather phenomena have real costs. Climate conditions also affect decisions concerning travel and leisure.

      As an exceptionally adaptable species, humans are found in all corners of the planet save the coldest polar regions, highest mountain peaks, and lowest ocean troughs. As a result, there remain few places on Earth that are not in some respect aptly classified as human-dominated ecosystems. No region is untouched by human influence. The release of waste products from domestic and economic enterprises (burning fossil fuels, synthetic chemical use, trash production, etc.) alters the composition of the atmosphere, and gases and particulates related to these activities travel to all parts of the globe. In contrast, land clearing and development often permanently alter the surface of the planet and modify patterns of both surface heating and local weather. Economically, these enterprises are purposeful actions; however, they are inadvertent when it comes to the realized changes in the environment. Air pollution (pollution), ozone depletion, acid precipitation, global warming, desertification, smog production, and deforestation are but a few of the human impacts on the climate system that arise from the alteration of the mass and energy exchange with the atmosphere.

      The Gaia hypothesis, introduced at the beginning of this section, remains controversial in the scientific community. Nevertheless, the question of whether Homo sapiens possesses the capacity and the will to maintain the Earth-atmosphere system in a sort of relative homeostatic balance is intriguing. The Montreal Protocol (United Nations Environment Programme) on Substances That Deplete the Ozone Layer, which has resulted in the phasing out of chlorofluorocarbons (chlorofluorocarbon) (CFCs), a group of industrial compounds that react with and disassociate ozone molecules, is a collective adaptive response by humans to a perceived and predicted threat to life from stratospheric ozone depletion. The Environmental Protection Acts ratified by the United Kingdom and Australia and the Kyoto Protocol to the United Nations Framework Convention on Climate Change are some examples of attempts to combat deleterious environmental change associated with the release of additional carbon dioxide into the air. If humans are to maintain the Earth-atmosphere system, it is through the social institutions of research and education that they will recognize environmental threats, come to wise decisions, and attempt to exercise stewardship of the planet.

Bruce P. Hayden

Additional Reading

General works
Introductory works include Edward Aguado and James E. Burt, Understanding Weather and Climate, 4th ed. (2007); P. Kabat et al. (eds.), Vegetation, Water, Humans, and the Climate: A New Perspective on an Interactive System (2004); Intergovernmental Panel on Climate Change: Working Group I, Climate Change 2007: The Physical Science Basis (2007); and National Research Council (U.S.), Panel on Climate Change Feedbacks, Understanding Climate Change Feedbacks (2003). Two excellent comprehensive reference works are John E. Oliver (ed.), The Encyclopedia of World Climatology (2005); and Stephen H. Schneider (ed.), Encyclopedia of Climate and Weather, 2 vol. (1996). Definitions of meteorological terms are provided in Todd S. Glickman (ed.), Glossary of Meteorology, 2nd ed. (2000); and Secretariat of the World Meteorological Organization, International Meteorological Vocabulary, 2nd ed. (1992), including nomenclature in English, French, Russian, and Spanish.Current research is reported in the following journals: Bulletin of the American Meteorological Society (monthly); Climatic Change (6/yr.); International Journal of Biometeorology (6/yr.); Journal of Applied Meteorology (monthly); International Journal of Climatology (15/yr.); Journal of Meteorological Research (bimonthly); International Journal of Meteorology (10/yr.); Monthly Weather Review; Quarterly Journal of the Royal Meteorological Society; Russian Meteorology and Hydrology (monthly); Global Climate Review (monthly); Weather (monthly); Weatherwise (bimonthly); and WMO Bulletin (quarterly).

Solar radiation and temperature
Introductory discussions of these basic elements of climate can be found in Grant W. Petty, A First Course in Atmospheric Radiation, 2nd ed. (2006); Craig F. Bohren and Eugene E. Clothiaux, Fundamentals of Atmospheric Radiation: An Introduction with 400 Problems (2006); and K.N. Liou, An Introduction to Atmospheric Radiation, 2nd ed. (2002).

Atmospheric humidity and precipitation
Discussions of water vapour in the atmosphere and global water budgets are found in Neil Wells, The Atmosphere and Ocean: A Physical Introduction, 2nd ed. (1997). Forms of precipitation are surveyed in D.M. Gray and D.H. Male (eds.), Handbook of Snow: Principles, Processes, Management & Use (2004). Eugenia Kalnay, Atmospheric Modeling, Data Assimilation, and Predictability (2003), provides a background on the statistics involved in numerical weather prediction. Technical treatments of cloud physics and development include R.R. Rogers and M.K. Yau, A Short Course in Cloud Physics, 3rd ed. (1989); and Robert A. Houze, Cloud Dynamics (1993), which also contains a pictorial cloud atlas.

Atmospheric pressure and wind
General textbooks with effective discussions of wind and pressure include Frederick K. Lutgens and Edward J. Tarbuck, The Atmosphere: An Introduction to Meteorology, 10th ed. (2007); and C. Donald Ahrens, Meteorology Today: An Introduction to Weather, Climate, and the Environment, 8th ed. (2007). The dynamics of atmospheric circulation and the use of general circulation models (GCMs) are documented in Masaki Satoh, Atmospheric Circulation Dynamics and Circulation Models (2004). More-sophisticated treatments of the wind-pressure relationship are provided in John A. Dutton, The Ceaseless Wind: An Introduction to the Theory of Atmospheric Motion, enlarged ed. (2002); James R. Holton, An Introduction to Dynamic Meteorology, 4th ed. (2004); and John M. Wallace and Peter V. Hobbs, Atmospheric Science: An Introductory Survey, 2nd ed. (2006).Roger A. Pielke, Sr.

Bin Wang, The Asian Monsoon (2006), is a wide-ranging work that summarizes the dynamics, climatic variability, forecasting, and modeling of the monsoon in Asia. Other introductory appraisals concerning monsoon development and behaviour are found in Roger A. Pielke, Jr., and Roger Pielke, Sr., Storms, 2 vol. (2000); Alexander Frater, Chasing the Monsoon: A Modern Pilgrimage Through India (2005); and C.P. Chang and T.N. Krishnamurti, Monsoon Meteorology (1987). A more technical consideration of monsoons that includes reviews of recent scientific investigations is found in C.P. Chang (ed.), East Asian Monsoon (2004).T.N. Krishnamurti

Upper-level winds
General works describing the generation and movement of upper-level circulation include Jonathan E. Martin, Mid-latitude Atmospheric Dynamics: A First Course (2006); James R. Holton, An Introduction to Dynamic Meteorology, 4th ed. (2004); and Karin G. Labitzke and Harry van Loon, The Stratosphere: Phenomena, History, and Relevance (1999). The relationship between upper-level circulation and weather forecasting is treated in Howard B. Bluestein, Synoptic-Dynamic Meteorology in Midlatitudes, 2 vol. (1993). Toby N. Carlson, Mid-latitude Weather Systems (1991), is an eloquent and comprehensive, but unassuming, account that explores the technical side of synoptic forecasting.Howard B. Bluestein

Climatic classification
Useful introductory discussions on the classification of climatic zones can be found in Alan H. Strahler and Arthur N. Strahler, Introducing Physical Geography, 4th ed. (2006); and Tom L. McKnight and Darrel Hess, Physical Geography: A Landscape Appreciation, 9th ed. (2007). Advanced treatments are provided in P.J. Robinson and A. Henderson-Sellers, Contemporary Climatology, 2nd ed. (1999); and Roger G. Barry and Richard J. Chorley, Atmosphere, Weather, and Climate, 8th ed. (2003). The global distribution of major climate types is the subject of Glenn R. McGregor and S. Nieuwolt, Tropical Climatology: An Introduction to the Climates of the Low Latitudes, 2nd ed. (1998). The foundational work on modern climate classification is contained in Glenn T. Trewartha and Lyle H. Horn, An Introduction to Climate, 5th ed. (1980). Particular regions are examined in H.E. Landsberg (ed.), World Survey of Climatology, 16 vol. in 18 (1969–2001); and the oft-celebrated account contained in Glenn T. Trewartha, The Earth's Problem Climates, 2nd ed. (1981).A. John Arnfield

Climate and life
A brief history of the coevolution of life and the atmosphere combined with the environmental consequences of changing the composition of the modern atmosphere is presented in Karl K. Turekian, Global Environmental Change: Past, Present, and Future (1996). Additionally, the work by William R. Cotton and Roger A. Pielke, Sr., Human Impacts on Weather and Climate, 2nd ed. (2007), chronicles the effects of human-induced landscape modification upon the dynamics of regional weather and climate.Interactions between organisms, the ecological systems they inhabit, and the atmosphere are detailed in William P. Lowry and Porter P. Lowry, Fundamentals of Biometeorology: Interactions of Organisms and the Atmosphere, 2 vol. (2001); and Stephen H. Schneider, Encyclopedia of Climate and Weather, 2 vol. (1996). The Gaia hypothesis is outlined in James E. Lovelock, Gaia: A New Look at Life on Earth (2000); and the role played by the biosphere in controlling the atmosphere and the oceans throughout geologic time is explained in James E. Lovelock, The Ages of Gaia: A Biography of Our Living Earth, rev. ed. (1995). Causes of climatic change are explained in detail in Bert Bolin et al. (eds.), The Greenhouse Effect, Climatic Change, and Ecosystems (1989). The first systematic treatment of the chemical and physical significance of atmospheric trace gases produced by the biosphere, including most of the early work on the greenhouse effect, is performed in John Tyndall, Heat Considered As a Mode of Motion (1871).Bruce P. Hayden

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