The atmosphere is the nearly transparent envelope of gases and suspended particles that surrounds the Earth, profoundly influencing environmental conditions on the planet's surface. Without chemical processes involving several of the atmospheric gases, life could not exist. The physical processes that operate in the atmosphere are also of vital importance because they are responsible for the Earth's varied climates. History of Studies Wind vanes were known to the ancient Chinese and Egyptians, and an atmospheric observatory, the Tower of the Winds, was built by the Greeks in the 2d century BC. Aristotle's interpretations of atmospheric phenomena in his Meteorologica, written about 340 BC, dominated Western thinking for almost 2,000 years (until the end of the 16th century). Modern study of the atmosphere began with the invention of the thermometer (c.1600) by \TGalileo\t and of the barometer (1643) by \TTorricelli\t. Early scientists quickly discovered that weather conditions are related to pressure changes, and by the end of the 17th century, records of atmospheric conditions were being kept at many locations. Using such records, Benjamin \TFranklin\t first deduced (1743) that storms are traveling systems, but this discovery could not result in useful weather forecasts until the telegraph was invented in the 1850s. During the latter half of the 19th century, \Lballoon\ls were used to probe the upper levels of the atmosphere. These efforts culminated in the discovery (1899) of the \Tstratosphere\t by a French meteorologist, Teisserenc de Bort. With the development in 1927 of the radiosonde (an inexpensive balloon-borne instrument package radioing back to Earth information on upper-level temperature, pressure, and humidity), systematic mapping of the atmosphere's structure and circulation began. At about the same time the theory that cyclonic storms arise from disturbances in polar fronts was developed by the Bergen school of meteorologists in Norway. Further key theoretical advances were made during the 1930s and '40s by Carl-Gustaf Rossby, a Swedish-American meteorologist who clarified the role of traveling weather disturbances in the circulation of the atmosphere as a whole. During World War II, bomber pilots discovered \Ljet stream\ls, narrow "rivers" of high-speed winds encircling the globe at altitudes of approximately 10 km (6 mi). With the launching (1960) of the first weather satellite, Tiros I, scientists were able for the first time to obtain a truly global view of the atmosphere. The concurrent development of high-speed computers has made it possible to construct mathematical models of complex atmospheric processes and to forecast the weather with increasing accuracy by numerical methods. COMPOSITION AND STRUCTURE Many of the physical and chemical processes that occur in the atmosphere are directly related to its composition. The atmosphere is now composed almost entirely of oxygen and nitrogen in their diatomic forms (two atoms bound together by chemical forces). Diatomic nitrogen accounts for approximately 78% of the total molecules in the atmosphere, and diatomic oxygen represents nearly 21%. The inert noble gas, argon, accounts for about 0.9%, and the remaining 0.1% is composed of many trace gases, the most significant of which are carbon dioxide and water vapor. Although carbon dioxide makes up only 325 parts per million of the atmosphere by volume, it is vital in maintaining the Earth-atmosphere system's heat balance because it strongly absorbs infrared (thermal) radiation. Water vapor, which is present in highly variable quantities ranging from 0 to 4% by volume, also absorbs considerable infrared radiation and, additionally, is an essential link in the \Thydrologic cycle\t. Another important trace gas is the triatomic form of oxygen, ozone, which is concentrated in a layer centered at about 25 km (16 mi) above the surface. Although present in maximum concentrations of only about 12 parts per million, ozone absorbs radiation in the ultraviolet region of the spectrum so effectively that the \Tozone layer\t can almost completely shield life on Earth from harmful ultraviolet rays. Meteorologists usually divide the atmosphere into four layers. In order of increasing elevation these are the \Ttroposphere\t, the stratosphere, the \Tmesosphere\t, and the \Tthermosphere\t. Each has a different temperature range. Temperatures decrease with altitude in the troposphere and mesosphere and increase with altitude in the stratosphere and thermosphere. The troposphere and stratosphere are separated by the \Ttropopause\t, a level of minimum temperature that varies in altitude from about 16 km (10 mi) near the equator to 9 km (5 mi) near the poles. The stratosphere and mesosphere are separated by the stratopause, a level of temperature maximum at an altitude near 50 km (30 mi). The mesosphere and thermosphere are in turn separated by a temperature minimum, the mesopause, which occurs near 80 km (50 mi). These temperature layers are created primarily by the selective absorption of \Tsolar radiation\t at various levels in the atmosphere. Radiation in the extreme ultraviolet (wavelength less than 100 nanometers) is absorbed by atoms of oxygen above 100 km (60 mi). This process not only maintains the high temperatures of the thermosphere but also produces electrically charged particles, called ions. For this reason the region of the atmosphere above 80 km (50 mi) is also referred to as the \Tionosphere\t. Ultraviolet radiation of somewhat longer wavelengths (200-300 nanometers) penetrates into the stratosphere, where it is absorbed by ozone to produce the temperature maximum near 50 km (30 mi). Visible radiation, on the other hand, penetrates to the surface of the Earth and produces the temperature maximum at the ground level. Stratification and Static Stability Temperature distribution is not alone in determining the state of the atmosphere. Pressure and density also are important. Atmospheric pressure, usually expressed in units called millibars, is the force that the total mass of air in an imaginary vertical column exerts on a given horizontal area of the Earth's surface. Standard sea-level pressure, 1,013.25 millibars, is equivalent to the pressure exerted by a column of mercury 760 mm (30 in) high. If, like water, the atmosphere were incompressible, pressure would decrease uniformly with height, and the atmosphere, like the ocean, would have a definite upper limit. In reality the atmosphere is compressible; that is, density (mass per unit volume) is proportional to pressure. This relationship, called \TBoyle's law\t, implies that density decreases with height in the atmosphere: as height increases, less mass remains above a given point; therefore less pressure is exerted. At sea level the density of air is about 1 kg per cu m (8 oz per cu ft). Both pressure and density decrease by about a factor of 10 for every 16 km (10 mi) increase in altitude. Density does not depend only on pressure; for a given pressure, it is inversely proportional to temperature. This relationship, known as Charles's law, implies that the depth of an air column bounded by two constant-pressure surfaces will increase as the temperature in the column increases. Thus the vertical distance over which pressure decreases to half of its surface value ranges from about 5,800 m (19,850 ft) in the tropics to 5,100 m (16,575 ft) near the poles. When an \Tair mass\t rises, it expands (because of the reduction in pressure). In expanding, it must work against the pressure force exerted by the surrounding air. According to the principle of conservation of energy (the first law of thermodynamics), the work done in expansion must be balanced by an equal reduction in the internal energy of the air mass. Since the internal energy is proportional to temperature, an expanding air mass must cool. Conversely, an air mass that is compressed must warm. Temperature changes caused by compression or expansion of a gas in the absence of heat exchange with the surroundings are called adiabatic changes. The process of adiabatic cooling or heating is essential to an understanding of vertical convection in the atmosphere. A mass of dry air rising adiabatically in the atmosphere cools at a rate of about 10 deg C per km (17 deg F per mi). Since, on the average, the decrease of temperature with height (the \Tlapse rate\t) in the troposphere is only 6.5 deg C per km (11 deg F per mi), an adiabatically ascending air mass becomes cooler and denser than its surroundings and tends to sink back toward its original level. Thus the mean vertical temperature structure is said to be statically stable with respect to a dry adiabatic displacement. In regions such as \Ldesert\ls, where the air near the ground is strongly heated, the lapse rate of temperature in the lower troposphere may exceed 10 deg C per km (17 deg F per mi). In such situations an air mass displaced vertically upward will become warmer than its surroundings and will be accelerated farther upward. Such statically unstable conditions result in vigorous mixing and upward heat transport, which tends to reduce the lapse rate toward the average value observed elsewhere. Moist Processes Water can exist in all three chemical phases--solid, liquid, and gas--at atmospheric temperatures and pressures. During phase changes, water exchanges heat with its surroundings, a process called latent heating. Condensation and freezing release heat and thus warm the surroundings, whereas evaporation and melting absorb heat and cool the surroundings. \TClouds\t and \Tprecipitation\t are thus essential to the heat balance of the atmosphere. The amount of water vapor present in a region of the atmosphere is usually expressed in terms of the ratio of the mass of water vapor in a given volume of air to the mass of that volume of dry air (the mixing ratio), the units used being grams of water per kilogram of dry air. Typically, this ratio is a few grams per kilogram (g/kg), but it can reach the range of 20 g/kg in the tropics, since warm air can hold more water than cold air. The value of the mixing ratio for saturated air--air that holds all the water it can at a given temperature--is called the saturated mixing ratio. Another common measure of water-vapor content is relative \Thumidity\t, defined as the ratio of the actual mixing ratio to the saturation mixing ration at the same temperature. Because the saturation mixing ratio decreases rapidly with decreasing temperature, when an air mass containing water vapor is forced to rise and cool adiabatically, it will eventually reach saturation. If it rises beyond this point, condensation will occur and clouds form. Condensation causes release of latent heat, which partially offsets adiabatic cooling. The net temperature decrease in an adiabatically expanding mass of saturated air will be less than that in a mass of dry air. This saturated-adiabatic lapse rate varies with temperature but is about 6 deg C/km in the midtroposphere. Thus, unlike a dry air mass, an ascending saturated air mass tends to become warmer and less dense than its environment, and thus to rise farther. The atmosphere is then said to be conditionally unstable because, although it is stable with respect to dry adiabatic ascent, it is unstable with respect to saturated adiabatic ascent. Conditional instability is a widespread occurrence and accounts for the predominance of convective (\Tcumulus\t) clouds in the atmosphere. The resulting vertical mixing plays an important role in maintaining the atmosphere's heat and momentum balances. Radiation and Energy Transfer In the long-term average, the energy content of the Earth-atmosphere system is nearly constant: energy received from the Sun is almost exactly balanced by energy radiated to space by the atmosphere. The average flux of solar energy incident of the top of the atmosphere per unit area perpendicular to the Sun's rays is about 1,380 watts/sq m (see \Tsolar constant\t). Not all of this energy is absorbed, however. A fraction is reflected back to space by clouds and by the Earth's surface. This fraction, called the \Talbedo\t, is about 30%. About 19% of the incident energy is absorbed in the atmosphere. The remaining 51% is absorbed at the Earth's surface (\Tinsolation\t). The sum of the energy absorbed by the Earth and the atmosphere must be balanced by thermal energy radiated into space. The amount of energy radiated by a body depends very much on its temperature. According to the Stefan-Boltzmann law, a body that radiates at the maximum possible efficiency (a so-called blackbody) emits an amount of energy that is proportional to the fourth power of the temperature. This law implies that the Earth balances the absorbed solar radiation by emitting radiation to space at an effective planetary temperature of -20 deg C (-4 deg F). This is much colder, however, than the average +15 deg C (+59 deg F) temperature of the Earth's surface. The excess surface temperature is maintained by the atmospheric \Tgreenhouse effect\t. Much of the infrared radiation emitted by the Earth is absorbed by water vapor and carbon dioxide in the atmosphere. The atmosphere reemits this radiation, both upward and downward. The fraction emitted downward is absorbed by the Earth's surface, thereby providing an energy source to maintain the surface temperature above the equilibrium value that would occur without an atmosphere. The intensity of the effect depends on the concentration of absorbing gases in the atmosphere. The radiation emitted by the atmosphere exceeds the amount absorbed from both solar radiation and infrared radiation from the ground by an amount equal to 30% of the incident solar flux. This energy deficit is balanced by heat transferred to the atmosphere from the Earth's surface primarily through the release of latent heat by evaporation at the surface and subsequent condensation in clouds. The annual average solar radiation received by the atmosphere varies strongly with latitude--it is four times greater at the equator than at the poles. Radiation emitted to space by the Earth, however, does not vary greatly from one latitude to another. This results in a net radiative energy surplus in the tropics and a net deficit in high latitudes. Thus, in order to maintain the Earth's overall energy balance, energy must be transported poleward. Observations indicate that poleward transport of heat in warm ocean currents, such as the \TGulf Stream\t, accounts for 50% of the required transport. The remaining 50% is transported by wind systems. CIRCULATION The average distribution of WINDS in the atmosphere is called the general circulation. The general circulation includes not only the more or less steady, global-scale winds but also the many transient disturbances that constitute the weather. To understand the structure of the wind and pressure systems on a global scale, the effects of the Earth's rotation on moving air parcels must be considered. On a nonrotating planet, differences in air pressure would be the primary cause of air flow. Winds would blow from high-pressure to low-pressure areas. The atmosphere is warmer at the equator, and a mass of warm air occupies more space than an equivalent mass of cold air; thus, at higher altitudes more air is displaced vertically upward at the equator than at the poles, which means that air pressure is greater at the equator at those altitudes. At higher altitudes on a nonrotating Earth, air would simply flow polewards. This would cause a buildup of sea-level air pressure at the poles, producing an equatorward flow of air at lower altitudes. The Earth's rotation, however, causes the flow of air to be deflected to the right of its direction of motion in the Northern Hemisphere and to the left in the Southern. This is called the \TCoriolis effect\t (after the 19th-century French mathematician G. G. Coriolis), and it is a result of the conservation of momentum. An air flow slows down or speeds up according to whether it moves farther from (equatorward) or nearer to (poleward) the Earth's axis of rotation. The pattern of air flow then becomes such that flow is parallel to isobars, or lines of constant air pressure at a given height, and the speed of flow is proportional to the pressure gradient. The wind is then said to be in geostrophic balance. The outcome is stated in \TBuys Ballot\t's law: if one stands in the Northern Hemisphere with one's back to the wind, air pressure is lower on the left than on the right. Temperature differences also affect the general pattern of air flow. The cooler the atmosphere is, the more rapidly air pressure decreases with height above the Earth's surface. Therefore the difference in pressure at two different latitudes also increases with height, and wind velocity must increase with height as well, if it is to maintain geostrophic balance. This tendency is described by the thermal-wind relationship, which states that if one stands with the warmer air to one's right and colder air to one's left in the Northern Hemisphere, the wind increases with height in the direction one faces. The geostrophic and thermal-wind relationships account for the mean winds being westerly (from the west) in mid-latitudes, where both pressure and temperature decrease poleward, and also account for the winds increasing with height to a maximum of about 30 m/sec in the subtropical jet stream core at an average of 30 deg latitude and 12 km elevation. These upper-level \Twesterlies\t are among the most important phenomena of the atmospheric general circulation. Transient areas of fair and bad weather originate in, and are carried along by, these upper-level winds. Thus, storm systems tend to move from west to east in mid-latitudes along storm tracks associated with the jet stream. The average jet stream winds do not, however, blow uniformly in the west-to-east direction; otherwise they could not transport heat poleward. In reality, the disturbing influence of large mountain ranges and land-sea contrasts (see \Tocean-atmosphere interaction\t) causes the jet stream to deviate from its mean latitude along a sinuous path circling the globe. In the Northern Hemisphere the largest equatorward deviations occur over the eastern parts of the North American and Asian continents. East of these regions, in the western Pacific and western Atlantic, the mean flow has strong poleward components that enable it to transport considerable heat northward. The same regions also create many transient weather disturbances, which move along with the mean winds and also act to transport heat poleward. In the tropical regions a different mechanism must be considered. Near the surface, the pressure decreases from the subtropical highs centered at approximately 30 deg latitude to the equatorial trough, or low-pressure zone (see \Ttropical climate\t). The resulting low-level winds depart from geostrophic balance because of the drag exerted by surface friction, and they form the northeasterly \Ttrade winds\t in the Northern Hemisphere and southeasterly trade winds in the Southern Hemisphere. The trade winds form the low-level branch of a closed-circulation pattern in the meridional plane known as the \THadley cell\t. The Hadley cell is named for the 18th-century Englishman George Hadley, who first proposed this circulation pattern. According to Hadley, surface air heated near the equator rises and moves poleward in the upper troposphere, gradually cooling and sinking back to the surface, where it again moves equatorward to close the loop. The actual Hadley circulation is highly asymmetric. The upward motion is concentrated along a narrow line about 50 km (30 mi) wide called the intertropical convergence zone, which, on the average, circles the globe about 10 deg north of the equator. Nearly all the upward mass transport in this zone occurs in the cores of intense cumulus convective clouds, which often extend to heights of 15 km (9 mi). Outside this zone, air gradually sinks as it is radiatively cooled over most of the subtropical belt. This widespread sinking accounts for the predominantly dry conditions in the subtropical zone. The heat and momentum transported out of the equatorial zone by the upper-level poleward flow in the Hadley cell are primarily responsible for maintaining atmospheric heat and momentum balances in the tropical zone. Weather Disturbances The transient disturbances that occur in the upper-level jet stream are responsible for much of the weather variations in extratropical latitudes. Small and random perturbations introduced into the jet-stream flow are amplified by drawing energy from the main flow. The resulting disturbances, typically several thousand kilometers in horizontal extent, are the familiar \Lcyclone\ls AND ANTICYCLONES seen on the daily weather charts. These systems go through life cycles of growth and decay that last about a week. They tend to concentrate preexisting temperature contrasts into narrow zones of sharp temperature change known as \Lfront\ls. Much of the precipitation associated with extratropical cyclonic systems is concentrated along the warm front (a transition zone from cool, dry air to the warm, moist air pushing it along) that precedes the arrival of the low-pressure minimum at the surface. The low is followed at the surface by passage of a cold front, which pushes out the warm air and marks a return to cool, dry conditions. (See \Tmeteorology\t.) Cyclonic storm systems also occur in the tropics, especially in association with the intertropical convergence zone. These systems appear as wavelike perturbations in the wind and as cloudiness fields that extend several thousand kilometers. Tropical disturbances, unlike mid-latitude systems, move from east to west, typically at a rate of about 10 deg longitude per day. Wind speeds associated with such disturbances ordinarily do not exceed several meters per second, although precipitation may exceed 2 cm/day (0.8 in/day). Occasionally, such disturbances intensify into tropical storms, or hurricanes, especially in the western Atlantic and western Pacific, where sea-surface temperatures are high (see \Thurricane and typhoon\t). Many of the most damaging storms are severe small-scale storms 10-100 km (6-60 mi) in horizontal extent, rather than the cyclonic variety. Isolated \Lthunderstorm\ls or squalls (see \Tsquall and squall line\t) generated by vertical motions produced by the heating of surface air are common during the summer months in areas where sufficient moisture is present and the atmosphere is conditionally unstable. When such systems become organized into squall lines, they may be accompanied by high winds, hailstorms, and even \Ltornado\les. These small-scale disturbances are fairly well understood by now, but little is yet known about the fundamental causes of climatic change (see \Tclimate\t), which occurs on all time scales. Long-term changes may be related to changing external forces such as variations in the Sun's energy output (see \Tsolar constant\t), but short-term variations seem to be random natural fluctuations in the atmosphere-ocean system. Human Impacts Events in recent years have created a greatly increased awareness of human impact on the atmosphere and the impact of the atmosphere on humans. The best-known example of deliberate human impact is the use of cloud-seeding techniques to increase precipitation and prevent hail. (See \Tweather modification\t.) Despite many years of extensive efforts, cloud seeding remains highly controversial, with very few well-substantiated successes. Inadvertent human modification of the atmosphere may have both local and global consequences. The local climatic impact of urbanization is well known (see \Turban climate\t). Global impacts are harder to assess. In recent years concern has grown over the possible damage to the ozone layer that might occur due to ozone-destroying chemical reactions with \Lfluorocarbon\ls introduced by human activities (see \Tpollution, environmental\t). In 1988 the major manufacturers of fluorocarbons agreed to phase out the production of these harmful chemicals. Some scientists estimate that each 1 percent reduction of the ozone layer by fluorocarbons would increase the amount of ultraviolet solar radiation reaching the Earth's surface by 1 to 3 percent. This would increase the incidence of cancer and could eventually threaten all life forms; it would also change the temperature distribution in the stratosphere, with potential global climatic effects. Scientists established in 1986 that an "ozone hole" (area of ozone depletion) develops every spring above Antarctica, and evidence gathered since then suggests that a similar phenomenon occurs over the Arctic. Interpreting the significance of this decline is difficult because of the complexity of atmospheric chemistry, but fluorocarbons are now considered to be primarily responsible for the decrease. At the Earth's surface, ozone levels are increasing as a result of air pollution. Tropospheric ozone damages crops and plays a role in \Tacid rain\t formation. The release of carbon dioxide by the burning of fossil fuels has been projected as another serious hazard if it continues unchecked. Concentrations of the gas in the atmosphere have risen from less than 300 parts per million (ppm) before 1900 to about 350 ppm in the late 1980s. If this trend continues into the next century, the result could be a substantial increase in global surface temperature because of the \Tgreenhouse effect\t. This would lead to regional climate changes and perhaps even to melting of the polar ice caps. Some scientists think that processes such as increased absorption of carbon dioxide by the ocean or an increase in cloud cover could offset the warming trend. Reports from the U.S. Environmental Protection Agency and the National Aeronautic and Space Administration, however, suggest that the trend is firmly established, and further research substantiates this view. Scientists have also found atmospheric methane to be increasing at a rate of about 1 percent each year as a result of farming practices, further depleting the ozone layer and enhancing the greenhouse effect. J. R. Holton THE ORIGIN OF THE ATMOSPHERE To understand how the atmosphere formed, scientists must examine 4 1/2 billion years of \Tgeologic time\t. Then, principles from chemistry and biology must be applied to develop a consistent explanation of the changes that must have occurred. Scientists generally agree that the Earth's atmosphere formed from gases emitted by \Lvolcano\les. Here, a problem arises, however. The composition of volcanic gases today is radically different from the composition of the atmosphere; in particular, volcanoes emit virtually no oxygen. If, as some evidence suggests, the gases emitted by volcanoes in the Earth's early years were the same as now, how did the atmosphere change so drastically? Volcanic gases in the young planet would have undergone a number of physical and chemical changes upon leaving the hot, pressurized interior of the Earth. Upon cooling, most of the water vapor would have condensed, filling the oceans. Much of the light hydrogen would then have escaped the gravitational attraction of the Earth. The carbon dioxide would have reacted with surface minerals, producing \Tcarbonate minerals\t and rocks. None of the changes, however, would have produced the oxygen essential for higher forms of life. Considerable evidence exists for the absence of oxygen during the first billion years of the Earth's existence. First, the earliest materials show incomplete oxidation. The Blind River uranium deposits of Canada, for example, contain uraninite, a \Turanium mineral\t that decomposes when exposed to today's oxygen-rich air. Second, no known source of free oxygen exists. Finally, the generally accepted theories of the origin of life indicate that life formed in an absence of oxygen. Two theories have been proposed to explain how oxygen was produced. One is the breakdown of water vapor by ultraviolet light (photodissociation). This would produce free hydrogen and oxygen, according to the reaction: 2H(2)O + ultraviolet light = 2H(2) + O(2) However, because large quantities of hydrogen would have to escape Earth's gravity for this reaction to be effective--an unlikely possibility--photodissociation does not appear to have been the major source of the Earth's oxygen. A second, more likely source is life itself, mainly through the process called \Tphotosynthesis\t, in which carbon dioxide and water combine to produce carbohydrates and oxygen: 6CO(2) + 6H(2)O = C(6)H(12)O(6) + 6O(2) According to some estimates, approximately 99% of the total amount of free oxygen added to the atmosphere since the Earth's beginnings was produced by photosynthesis, with only 1% produced by photodissociation. If life produced most of the oxygen in the atmosphere, however, life itself must have originated in an environment that would have been extremely hostile to present forms of life. Without oxygen, the Sun's ultraviolet radiation, deadly to all cells, would reach the Earth's surface in lethal doses. Thus the first organisms must have had a very restricted habitat: under a layer of water, near enough to the surface to obtain sunlight yet deep enough to avoid the ultraviolet radiation. In an atmosphere without oxygen, lethal radiation penetrates to a depth of about 10 m (30 ft), and early forms of life probably lived near or slightly below this level, secreting oxygen as a fermentation product that, to them, was a dangerous poison. As these organisms lived, died, and sank to the bottom, oxygen and ozone slowly increased in the atmosphere. The ozone reduced the amount of ultraviolet radiation reaching the surface, making the upper layers of the ocean more hospitable. The organisms then invaded the layer of water immediately below the surface, where sunlight is abundant. They evolved to forms that could obtain energy through respiration as well as fermentation, and the process of photosynthesis began to accelerate. An explosive diversification of species occurred at this point, apparently within the Cambrian Period that began about 570 million years ago. Finally, the atmosphere's oxygen content increased enough to allow life to emerge from the water onto dry land, beginning in the Silurian Period about 430 million years ago. Presumably the oxygen content of the atmosphere has varied with major climatic changes of the past and with associated variations in photosynthesis, but such large-scale changes cannot be measured. On a smaller scale, scientists have found that oxygen is escaping into space from the polar regions of the ionosphere at a rate of about 5 x (10 to the 7th power) kg per year, but this loss is counterbalanced by the dissociation of water vapor. Richard A. Anthes Bibliography: Allen, Oliver, Atmosphere (1983); Anthes, Richard, et al., The Atmosphere, 3d ed. (1981); Bohren, Craig F., Clouds in a Glass of Beer (1987); Goody, R. M., and Walker, J. C. G., Atmospheres (1972); Gribben, John, The Breathing Planet (1986); Lutgens, F. K., and Tarbuck, E. J., The Atmosphere: An Introduction to Meteorology, 3d ed. (1986); Riehl, Herbert, Introduction to the Atmosphere, 2d ed. (1978); Walker, James C. G., Evolution of the Atmosphere (1977); Wallace, J. M., and Hobbs, P. V., Atmospheric Science: An Introductory Survey (1977); Wayne, Richard, The Chemistry of Atmospheres (1985). See also: \TEarth\t; \Tconservation, laws of\t; \Tgas laws\t; \Tweather forecasting\t.