ATMOSPHERE

gaseous envelope that surrounds the Earth. Other planets of the solar system, as well as a few of the large satellites of the outer planets, such as Saturn's Titan, also have atmospheres. The atmosphere that surrounds the Earth consists of a mixture of gases, primarily nitrogen and oxygen. This envelope, commonly called the air, also contains numerous less abundant gases, water vapour, and minute solid and liquid particles in suspension. Rocket probes and especially the drag encountered by artificial satellites at altitudes of several thousand kilometres have demonstrated that the terrestrial atmosphere extends to a very great distance. The composition of the atmosphere encodes a great deal of information bearing on its origin. Furthermore, the nature and variations of the minor components reveal extensive interactions between the atmosphere, terrestrial environment, and biota. The development of the atmosphere and such interactions are discussed in the first major section of this article, with particular attention given to the rise of biologically produced molecular oxygen, O2, as a major component of air. The atmosphere is considered in terms of layers, or regions, arranged like spherical shells above the surface of the Earth. The chemical and physical properties of these various regions are treated in considerable detail. Such upper atmospheric phenomena as airglow, auroras, and the Van Allen radiation belts are included in the coverage. gaseous envelope that surrounds the Earth and other planets of the solar system. A few of the large satellites of the outer planets, such as Saturn's Titan, also are known to have atmospheres. Near the Earth's surface, the atmosphere has a well-defined chemical composition, consisting of molecular nitrogen (78 percent by volume), molecular oxygen (21 percent), and argon (0.9 percent). It also contains small amounts of carbon dioxide and water vapour, along with trace quantities of methane, ammonia, nitrous oxide, hydrogen sulfide, helium, neon, krypton, xenon, and various other gases. Present too are solid and liquid particles in suspension. The density of the atmosphere at the terrestrial surface is 1.2 kg per cubic metre (0.075 pound per cubic foot), but it decreases rapidly with height. Atmospheric scientists commonly divide the atmosphere into five layers, or regions, on the basis of the vertical distribution of temperature. The troposphere, the region in contact with the Earth's surface and where weather occurs, is characterized by a decrease of temperature with increasing altitude. It extends in height from approximately 6 to 8 km (4 to 5 miles) at the poles to about 17 km (11 miles) at the equator. Above this region is the stratosphere, in which temperature rises with increasing altitude to about 50 km (30 miles). The troposphere and the stratosphere show distinct circulation systems. Whereas vertical motions prevail in the former, motions in the latter are largely confined to the horizontal. Above 50 km lies the mesosphere, which is characterized by a rapid decrease of temperature. Like the troposphere, this region is subject to strong seasonal variations of temperature at high latitudes. Beyond the mesosphere is the thermosphere, where temperature rapidly rises, attaining a maximum value of more than 1,000 C (1,832 F) at about 400 km (240 miles). The increase in temperature ceases at this height, beyond which lies the exosphere, the highest layer of the atmosphere. The density of the atmosphere is so low in this layer that molecular collisions rarely occur, and hence the concept of temperature loses its customary meaning. In the exosphere, light atoms, such as those of hydrogen and helium, may acquire sufficient velocity to escape the Earth's gravitational pull. The aforementioned mixture of molecular nitrogen and molecular oxygen prevails throughout the troposphere, stratosphere, and mesosphere, which together make up what is generally called the homosphere. However, in the stratospheric layer, solar ultraviolet radiation causes some molecular oxygen (O2) to dissociate, and the resulting atomic oxygen (O) then combines with the remaining O2 to form ozone (O3). The relatively high concentration of ozone in the stratosphere helps to prevent much of the Sun's ultraviolet radiation of shorter wavelength (which is harmful to plant and animal life) from reaching the Earth's surface. Volatile gases are constantly being added to the atmosphere by volcanic eruptions and hot springs. Far more significant than these natural sources, however, is industrial activity, particularly the combustion of fossil fuels (e.g., oil, natural gas, and coal). This process adds a variety of noxious gases such as carbon monoxide, hydrogen chloride, sulfur dioxide, and nitrogen oxide. The latter two play a central role in the formation of acid rain, which has become a major environmental problem. The burning of fossil fuels also releases large quantities of carbon dioxide to the atmosphere, which may affect the heat balance of the Earth in what has been called the greenhouse effect. Most of the atmosphere consists of neutral atoms and molecules, but at high altitudes a significant fraction is electrically charged owing to photoionization. This region, called the ionosphere, begins near the top of the stratosphere and extends essentially throughout the mesosphere and thermosphere but is most distinct at altitudes above roughly 80 km (50 miles). Another region of the atmosphere identified by parameters other than temperature gradient is the magnetosphere. This is a vast region in which charged particles have energies greater than those corresponding to thermal velocities and move along the flux lines of the Earth's magnetic field. also called Standard Atmosphere, unit of pressure, nearly equal to the mean atmospheric pressure naturally existing at sea level on the surface of the Earth or to the pressure exerted by a vertical column of mercury (as in a barometer) 760 mm (29.9213 inches) high, used in meteorology. It is defined as 101,325 pascals, or newtons of force per square metre (approximately 14.7 pounds per square inch). See also millibar. Additional reading Development of the Earth's atmosphere Astrophysical considerations bearing on the earliest stages of atmospheric development are reviewed in John S. Lewis and Ronald G. Prinn, Planets and Their Atmospheres: Origin and Evolution (1984). Aspects of subsequent development are discussed in James C.G. Walker, Evolution of the Atmosphere (1977); and Heinrich D. Holland, The Chemical Evolution of the Atmosphere and Oceans (1984). Numerous chapters in J. William Schopf (ed.), Earth's Earliest Biosphere: Its Origin and Evolution (1983), discuss biologic controls of atmospheric composition and their development over time. The probable rise of oxygen just prior to the development of the earliest animals is discussed in detail by Bruce Runnegar, The Cambrian Explosion: Animals or Fossils?, Journal of the Geological Society of Australia, 29(4):395411 (1982). Development of the biogeochemical cycle of carbon and its interactions with the atmosphere are discussed in E.T. Sundquist and W.S. Broecker (eds.), The Carbon Cycle and Atmospheric CO2: Natural Variations Archean to Present (1985). J.M. Hayes Structure, composition, and physical properties of the atmosphere John T. Houghton, The Physics of Atmospheres, 2nd ed. (1986), is a readable account of the physical basis for the treatment of atmospheric radiation and a good introduction to atmospheric dynamics. Simpler discussions are given by Richard M. Goody and James C.G. Walker, Atmospheres (1972). A useful account of atmospheric chemistry is given by Richard P. Wayne, Chemistry of Atmospheres: An Introduction to the Chemistry of the Atmospheres of Earth, the Planets, and Their Satellites (1985). Atmospheric composition, including an account of the interaction of the atmosphere with the oceans, is discussed by Heinrich D. Holland, The Chemistry of the Atmospheres and Oceans (1978). Effects of human activities on the atmosphere An overview of the Earth's endangered, changing atmosphere is given in John Gribbin (ed.), The Breathing Planet (1986), a collection of essays from New Scientist; and The Changing Atmosphere: Implications for Mankind, Chemical and Engineering News, vol. 64, no. 47 (Nov. 24, 1986), a special issue on global warming and the depletion of stratospheric ozone. For a review of current scientific understanding of these problems, see World Meteorological Organization, Atmospheric Ozone: An Assessment of Our Understanding of the Processes Controlling Its Present Distribution and Change, 3 vol. (1985), report no. l6 of the Global Ozone Research and Monitoring Project. See also William C. Clark (ed.), Carbon Dioxide Review, 1982 (1982); United States. Congress. Senate. Committee on Environment and Public Works. Subcommittee on Toxic Substances and Environmental Oversight, Global Warming (1986); and Harold W. Bernard, Jr., The Greenhouse Effect (1980). A survey of the history of acid deposition from the atmosphere and its sources is provided by Eville Gorham, Acid Rain: An Overview, ch. 1 in Chandrakant M. Bhumralkar (ed.), Meteorological Aspects of Acid Rain (1984), pp. 118, with an extensive bibliography. More detailed information may be found in Jon R. Luoma, Troubled Skies, Troubled Waters: The Story of Acid Rain (1984); and Thomas Pawlick, A Killing Rain: The Global Threat of Acid Precipitation (1984). Further bibliographic information is available in an annotated bibliography of research, G. Harry Stopp, Jr., Acid Rain (1985). The ionosphere and phenomena of the upper atmosphere For a general review of the underlying physics of the upper atmosphere, see J.A. Ratcliffe (ed.), Physics of the Upper Atmosphere (1960); and C.O. Hines et al. (eds.), Physics of the Earth's Upper Atmosphere (1965). A modern account of aeronomy is given in P.M. Banks and G. Kockarts, Aeronomy, 2 vol. (1973). A classic treatment of the aurora and airglow is presented in Joseph W. Chamberlain, Physics of the Aurora and Airglow (1961). For a specialized account of magnetospheric processes, see Wilmot N. Hess, The Radiation Belt and Magnetosphere (1968). A readable summary article is J.G. Roederer, The Particle and Field Environment of the Earth, Astronautics and Aeronautics, 7:2228 (January 1969).Current research is reported in the following journals: Advances in Atmospheric Sciences (quarterly); Atmospheric Environment (monthly); Environmental Science & Technology (monthly); Journal of the Atmospheric Sciences (semimonthly); Science (weekly); and Scientific American (monthly). Michael B. McElroy Composition of the present atmosphere Major components of the lower atmosphere The atmosphere contains a bewildering array of gases, with relative abundances for important species ranging from 78 percent (N2) to less than one part in 1012 (the hydroxyl radical, OH). The longer lived gasese.g., N2 and O2are distributed more or less homogeneously around the Earth. The shorter lived speciese.g., carbon monoxide, nitric oxide, and ozonecan vary considerably both in time and space. In many respects, the atmosphere can be considered an extension of the biosphere: almost all of the major constituents, with the exception of the noble gases, are either directly or indirectly under the influence of life. There are several natural linkages, as, for example, O2, CO2, CH4, and H2 (molecular hydrogen). Oxygen is a product of photosynthesis, summarized conveniently by the bulk reaction The notation is qualitative rather than quantitative. The formula CH2O denotes any of a variety of organic compounds formed in the primary life-giving photosynthetic event: the stoichiometry is approximately 1:2:1, C:H:O. Aerobic respiration and decay involve the reverse of reaction (9), This reaction satisfies the energy needs of the human population and of most of the other higher animals. All life on Earth ultimately depends on the ability of plants to capture solar energy and to store this energy in the form of potential food, CH2O. It is a two-way street. In the absence of reaction (10), carbon would accumulate in organic form and the fuel for photosynthesis, atmospheric CO2, would be depleted. Bacteria play a major role in recycling carbon primarily by reactions analogous to (10). Respiration and decay can proceed even when the supply of O2 is limited. This can arise, for example, in the sediments of organic-rich swamps and in the stomachs of ruminants. The product of anaerobic decay is methane (CH4) in this case. Oxidation of carbon may occur photochemically in the atmosphere, initiated by reaction with the OH radical: molecular hydrogen is a product of the oxidation of CH4 and other hydrocarbons; like CH4, H2 is removed from the atmosphere mainly by reaction with OH. Distribution of carbon, nitrogen, and oxygen compounds Carbon compounds The bulk of the Earth's volatile carbon resides in sediments, either as organic carbon or as a component of carbonate minerals such as calcite, CaCO3. Carbon is carried to the sediments in detrital material. The fraction lost from the oceans during subsequent burial represents but a small fraction of the total net primary productivityless than 1 percent. Figure 1: A schematic representation of the biogeochemical cycle of carbon. The life cycle is efficient (see above Figure 1). Carbon atoms are exchanged back and forth between the atmosphere, biosphere, soils, and oceans. Even the sediments provide only a temporary, albeit in human terms long (100,000,000-year), residence for the restive atom. The atom returns to the atmosphere as sediments are uplifted and weathered. The transit time from weathering to eventual return to the sediments is about 100,000 years, most of this spent as a component of the bicarbonate ion, HCO3-, in solution in the deep sea. On time scales of a few hundred years or longer, the abundance of CO2 in the atmosphere is determined by the dynamics, chemistry, and biology of the oceans. Most of the carbon in the oceans is present in cold, relatively stagnant water at depth. It returns to the atmosphere in association with slow upwelling motion at low latitudes. As surface waters cool and sink at high latitude, they draw carbon from the atmosphere, roughly balancing the source at low latitude. Falling fecal material provides an additional important means for transporting carbon from the surface to the deep. The dynamics of this complex exchange are just beginning to be understood. It is clear, though, that the level of atmospheric CO2 is not immutable. Studies of gases trapped in polar ice indicate that CO2 has fluctuated from about 200 to roughly 280 parts per million (ppm) over the past 100,000 years. Low levels of CO2 are associated with cold, ice-age conditions at the surface; high values correspond to times when the climate was relatively warm during interglacials. It appears that this behaviour has persisted over at least the past 750,000 years. It is against this background that assessments must be made of the impact of the recent change in the CO2 level caused by the burning of fossil fuels. The level of CO2 has risen since the Industrial Revolution from about 280 ppm in 1850 to approximately 350 ppm today. It is expected to climb to values in excess of 600 ppm by the early part of the 21st century. This poses a double challenge. Is it possible to predict accurately the effects of an increasing level of CO2 on climate? If so, can this knowledge be used to influence the course of action over the next few decades? Humankind has developed the ability to change the Earth on a global scale in a single lifetime. Yet, it remains to be seen whether scientific knowledge can develop apace. Structure of the present atmosphere General characteristics The atmosphere extends from the surface of the Earth to heights of thousands of kilometres, where it gradually merges with the solar wind. The composition of the atmosphere as measured by its mean density (the average mass per unit volume) is more or less constant with height to altitudes of about 100 kilometres. This state of approximate uniformity arises as a result of motion and as a consequence of the high frequency with which molecules of a particular species are involved in collisions with their neighbours. A representative oxygen molecule, O2, for example, encounters a nitrogen molecule, N2, on average once every 10-9 second at the surface. Even at heights of 100 kilometres, where the density of air molecules is much lower, the encounter time is still comparatively brief, about 10-3 second. A force imparted to one molecule is rapidly transferred to all. The atmosphere tends to behave as though it were composed of a single molecular species with an effective molecular mass set by its mean composition. The bulk of the lower atmosphere is composed of N2 and O2, with relative abundances of, respectively, 0.78 and 0.21, based on the average number of molecules present in a representative volume of air. The mass of the hypothetical mean molecule of the lower atmosphere is 28.97 atomic units (one atomic unit corresponds to the mass of a hydrogen atom, 1.66 10-24 gram). This value is intermediate between that of N2 (28 atomic units) and that of O2 (32 atomic units) and reflects the presence in the atmosphere of trace quantities of water (18 atomic units), argon (40 atomic units), carbon dioxide (44 atomic units), and other less abundant compounds as well. Figure 3: Average molecular mass of the atmosphere in atomic units (one atomic unit corresponds to The collisional interaction between individual molecules becomes progressively less efficient at altitudes above 100 kilometres. Molecules begin to experience a force of gravity proportional to their individual molecular masses. Heavy gases are bound more closely to the Earth, whereas lighter gases are free to float higher. The average molecular mass of the atmosphere therefore declines steadily with increasing altitude, as illustrated in Figure 3. Atomic oxygen is more abundant than N2 above about 160 kilometres. In turn, atomic oxygen gives way to helium above 600 kilometres, and hydrogen is the major constituent at altitudes higher than 1,000 kilometres. The region above 100 kilometres is referred to as the heterosphere, a name intended to emphasize the importance of the change in composition as a function of altitude. In the same vein, the region lower than 100 kilometres was given the name homosphere. Division based on thermal structure Figure 4: Thermal structure of the atmosphere showing depths of penetration for sunlight of A second classification, based on thermal structure, provides a more detailed and, in many respects, more useful scheme for the division of the atmosphere into distinct layers (Figure 4). The temperature decreases rapidly above the surface of the Earth to an altitude of about 17 kilometres. The air is relatively unstable, a consequence of the decrease of temperature with altitude. Warmer air is comparatively light and has a tendency to rise. Conversely, colder air is dense and tends to sink. The atmosphere is poised to turn over, to convect much like water in a kettle heated from below. This region is known as the troposphere, a term derived from the Greek words tropos, turning, and sphaira, ball. Most of the weather of the planet is confined to the troposphere. The upper boundary of the troposphere is called the tropopause. The temperature begins to increase slowly with altitude above the tropopause in a region known as the stratosphere, from the Latin word stratus, meaning stretched out or layered. Vertical motions are strongly inhibited in the stratosphere. An air parcel that attempts to rise becomes rapidly colder and denser than the air it displaces. Buoyancy forces in this environment act to suppress vertical motion. Motions in the stratosphere are thus largely confined to the horizontal, accounting for the layered structure of high-altitude stratus clouds. The increase of temperature with altitude persists to about 50 kilometres, at which point the temperature is about as high as at the surface. This marks the upper boundary of the stratosphere, the stratopause. The temperature resumes its general decrease with altitude above the stratopause in the mesosphere (mesos denoting middle). It reaches a minimum near 85 kilometres at the mesopause, which is the coldest region of the atmosphere. The temperature increases again with altitude above the mesopause in the thermosphere, so named because of the importance of thermal conduction in this region. A large portion of the heat deposited in the thermosphere is conducted downward and is radiated out to space from the vicinity of the mesopause. Figure 5: Energy budget for the surface of the Earth illustrating what happens, on average, to 100 The thermal structure of the atmosphere reflects in part the influence of energy deposited directly by the absorption of sunlight. It is determined, though, to a much larger extent by a complex suite of processes important to redistributing energy vertically. The Sun is the ultimate source of energy. Slightly more than 50 percent of the energy incident from the Sun is absorbed by the surface. A comparable amount, roughly 30 percent, is reflected back into space, either by clouds (20 percent), by air (6 percent), or by the surface itself (4 percent), as shown in Figure 5. The atmosphere absorbs only about 16 percent of the incident energy; most of this is captured by dust particles in the troposphere. The atmosphere is bathed in two more or less distinct radiation fields. The first field, originating in the Sun, has the majority of its energy in the visible and ultraviolet portions of the electromagnetic spectrum. The second, emanating from the surface of the Earth and its lower atmosphere, has most of its energy at longer wavelengthsnamely, in the infrared portion. Figure 6: Spectrum of the Sun compared with energy emitted by an ideal blackbody at 5,785 K. The solar spectrum at visible wavelengths is about what would be expected for a blackbody radiating at a temperature of 5,785 K, the temperature of the photosphere from which most of the solar radiation is emitted. (A blackbody is a hypothetical ideal body or surface that absorbs and reemits all radiant energy falling upon it.) Radiation at shorter wavelengths is more intense (see Figure 6). Light at ultraviolet and X-ray wavelengths emanates from the outermost regions of the solar atmosphere, the chromosphere and corona. Temperatures there climb to values above 106 K. Figure 7: Spectrum of the Earth as viewed from space showing distinction between reflected sunlight Viewed from space, the spectrum of the Earth would be similar to that shown schematically in Figure 7. At longer wavelengths the radiation would be emitted by the atmosphere and surface and derived more or less equally from the dayside and nightside of the planet. At shorter wavelengths the spectrum would be dominated by sunlight reflected by clouds and by the surface on the dayside. The ionosphere and phenomena of the upper atmosphere The ionosphere General characteristics The bulk of the atmosphere consists of electrically neutral atoms and molecules. At high altitudes, however, a significant fraction of the atmosphere is electrically charged. This region is generally called the ionosphere. It extends throughout the mesosphere and thermosphere but is most important and distinct at altitudes above about 80 kilometres. The name was introduced during the 1920s and was formally defined in 1950 by a committee of the Institute of Radio Engineers. Members of the committee identified the region as the part of the earth's upper atmosphere where ions and electrons are present in quantities sufficient to affect the propagation of radio waves. Much of the early research on the ionosphere was carried out by radio engineers and was stimulated by the need to define the factors influencing long-range radio communication. Priorities have changed in recent years. Today, the need is to understand the ionosphere as the environment for Earth-orbiting satellites and ballistic missiles. The emphasis is on processes. Scientific knowledge of the ionosphere has grown tremendously, fueled by a steady stream of data from spacecraft-borne instruments and enhanced by measurements of relevant atomic and molecular processes in the laboratory. Historically, the ionosphere was thought to be composed of a number of relatively distinct layers. The most important layers were identified by the letters D, E, and F, with the F layer subsequently divided into regions F1 and F2. The nomenclature is somewhat peculiar. It appears that Edward V. Appleton of Britain, a pioneer in early radio probing of the ionosphere, was accustomed to using the symbol E to describe the electric field of the wave reflected from the first layer of the ionosphere he studied. Later, Appleton identified a second layer at higher altitude and used the symbol F in this case. Suspecting a layer at lower altitude, he adopted the additional symbol D. In time the letters came to be associated with the layers themselves rather than with the field of the reflected waves. It is now known that Appleton's layers are not particularly distinct. The electron density increases more or less uniformly with altitude from the D region, reaching a maximum in the F2 region. In spite of the peculiarities of the original rationale for the naming of the layers, the nomenclature employed to describe the different regions of the ionosphere continues in wide use. Though the names persist, the definitions have evolved to reflect present-day understanding of the underlying physics and chemistry. Ionospheric physics and chemistry Most of the ionization in the ionosphere is effected by photoionization. Photons of short wavelength (i.e., high energy) are absorbed by atmospheric gases. A portion of the energy is used to eject an electron, converting a neutral atom or molecule to a pair of charged species: an electron, which is negatively charged, and a companion positive ion. Ionization in the F1 region is produced mainly by ejection of electrons from O2, O, and N2. The threshold for ionization of O2 corresponds to a wavelength of 102.7 nanometres. Thresholds for O and N2 are at 91.1 and 79.6 nanometres, respectively. Positive ions can react with neutral gases and change their identity. There is a tendency for these reactions to favour production of more stable ions. Thus, ionized oxygen, O+, can react with O2 and N2, resulting in ionized molecular oxygen, O2+, and ionized nitric oxide, NO+: Similarly, ionized molecular nitrogen, N2+, can react with O and O2, forming NO+ and O2+: The most stable, and consequently most abundant, ions in the E and F1 regions are O2+ and NO+, the latter more so than the former. At lower altitude, O2+ can be converted to NO+ by reactions with the minor species N (nitrogen) and NO (nitric oxide), In the D region, NO+ can be converted to H3O+, the hydronium ion, and to companion species such as H5O2+ and H7O4+, formed by the addition of H2O. Production of hydrated ions is limited by the availability of H2O. As a consequence, they are confined to altitudes below about 85 kilometres. The electron density in the D, E, and F1 regions reflects for the most part a local balance between production and loss. Electrons are removed mainly by dissociative recombination, a process in which electrons attach to positively charged molecular ions and form highly energetic, unstable neutral molecules. These molecules decompose spontaneously, converting internal energy to kinetic energy of fragments. The most important processes in the ionosphere involve recombination of O2+ and NO+. The reactions in this case may be summarized as follows: A portion of the energy released in reaction (67) may appear as internal excitation of either nitrogen, oxygen, or both. The excited atoms can radiate, emitting faint visible light in the green and red regions of the spectrum, and thereby contribute to the phenomenon of the airglow. The airglow originates principally from altitudes above 80 kilometres and is responsible for the diffuse background light that makes it possible to distinguish objects at the Earth's surface on dark, moonless nights. Airglow is produced for the most part by reactions involved in the recombination of molecular oxygen. The contribution from reaction (67) is readily detectable, however, and provides a useful technique with which to observe changes in the ionosphere from the ground. Studies of the airglow have a long and checkered history in atmospheric science. Over the years they have contributed significantly to scientific understanding of processes in the upper atmosphere. As indicated above, dissociative recombination provides an effective path for the removal of molecular ions. There is no comparable means for the removal of atomic ions. Direct recombination of O+ with an electron requires that the excess energy be radiated as light. Radiative recombination is inefficient, however, compared with dissociative recombination and plays a small role in the removal of ionospheric electrons. There is a complication at high altitudes where atomic oxygen is the major constituent of the neutral atmosphere and where electrons are produced primarily by photoionization of O. The atomic ion O+ may be converted to NO+ and O2+ through reactions with N2 and O2, but the abundances of N2 and O2 decline relative to O as a function of increasing altitude. In the absence of competing reactions, the concentration of O+ and the density of electrons would increase steadily with altitude, paralleling the rise in the relative abundance of O. This occurs to some extent but is limited eventually by vertical transport. Ions and electrons produced at high altitude are free to diffuse downward, guided by the Earth's magnetic field. The lifetime of O+ is long at high altitude where the densities of O2 and N2 are very small. As ions move downward, the densities of O2 and N2 increase. Eventually the time constant for reaction of O+ with O2 and N2 becomes comparable to the time for diffusion: O+ is converted to either O2+ or NO+ before it can move much farther. The O+ density exhibits a maximum in this region. Competition between chemistry and transport is responsible for the formation of an electron-density maximum in the F2 layer. The dominant positive ion is O+. The density of O+ decreases with decreasing altitude below the peak, reflecting a balance between production by photoionization of O and removal by reaction (64). The density of O+ also decreases above the peak. In this case, the removal of photo-ions is regulated by downward diffusion rather than by chemistry. The distribution of O+ with altitude above the peak reflects a balance of forces, a pressure gradient that acts to support O+ in opposition to gravitational and electrostatic forces that combine to pull O+ down. The electrostatic force is set up to preserve charge neutrality. In its absence, the density of ions, which are much more massive than electrons, would tend to fall off more rapidly with altitude than that of electrons. The abundance of electrons would quickly exceed the density of ions, and the atmosphere would accumulate negative charge. The electric field redresses the imbalance, drawing electrons down and providing additional upward support for positively charged ions. The abundance of O+ falls with increasing altitude as though O+ had a mass of 8 atomic units rather than 16 atomic units (the electric field exerts a force equivalent to the gravitational force on a body of mass 8 atomic units, directed upward for ions and downward for electrons). The density of electrons falls off with altitude at precisely the same rate as O+, preserving the balance of positive and negative charge. Ionization at any given level depends on three factors: (1) the availability of photons of a wavelength capable of effecting ionization, (2) a supply of atoms and molecules necessary to intercept this radiation, and (3) the efficiency with which the atoms and molecules are able to do so. The efficiency is relatively large for O, O2, and N2 from about 10 to 80 nm. This is the portion of the spectrum responsible for the production of electrons and ions in the F1 region. Photons between 90 and 100 nm are absorbed only by O2. They therefore penetrate deeper and are responsible for producing about half the ionization in the E layer. The balance is derived from soft X rays (those of longer wavelength), which are absorbed with relatively low efficiency in the F region and so are able to penetrate to altitudes of about 120 kilometres under conditions of high solar elevation. Hard X rays (those of shorter wavelength), notably below about 5 nm, reach even deeper. This portion of the spectrum accounts for the bulk of the ionization in the D region, with an additional contribution from wavelengths longer than 102.6 nmmainly from photons in the strong solar emission line at Lyman a, 121.6 nm. Lyman a emission is absorbed weakly by the major components of the atmosphere, O, O2, and N2. It is absorbed readily by NO and has sufficient energy to ionize this relatively unstable compound. In spite of the low abundance of NO, the high flux of solar radiation at Lyman a is able to provide a significant source of ionization for the D region near 90 kilometres.