atmosphere

Introduction
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.

Development of the Earth's atmosphere
A complete reconstruction of the origin and development of the atmosphere would include details of its size and composition at all times during the 4,500,000,000 years since the formation of the Earth. This goal could not be achieved without knowledge of the pathways and rates of supply and consumption of all atmospheric constituents at all times. Information regarding these processes, however, is incomplete even for the present atmosphere, and there is almost no direct evidence regarding atmospheric constituents and their rates of supply and consumption in the past.
The contrast with related fields of the Earth's history is notable. Fossils and other structural and chemical details of ancient rocks provide information useful to evolutionary biologists and historical geologists, but ancient atmospheres, “mere vapours,” have not left such substantial remnants. These vapours are, however, the stuff of stars and the moving force of storms and erosion. Although historians of the atmosphere must rely heavily on inference, the object of their interest has played so important a role in the Earth's history that evidence related to its development, though indirect, is abundant.

Concepts related to atmospheric development
The atmosphere as part of the crust
To the Earth scientist, the crust includes not only the top layer of solid material (soil and rocks to a depth of 6–70 kilometres, separated from the underlying mantle by differences in density and by susceptibility to surficial geologic processes) but also the hydrosphere (oceans, surface waters on land, and groundwater beneath the land surface) and the atmosphere. Interactions among these solid, liquid, and gaseous portions of the crust are so frequent and thorough that considering them separately introduces more complexities than it eliminates. As a result, a description of the history of the atmosphere must concern itself with all volatile components of the crust.

Materials
Volatile compounds as well as elements important in present and past atmospheres or in interactions between the atmosphere, biosphere, and other portions of the crust include the following:
Present major components: molecular nitrogen (N2) and molecular oxygen (O2)
Noble gases: helium (He), neon (Ne), argon (Ar), krypton (Kr), and xenon (Xe)
Abundant variable components: water vapour (H2O) and carbon dioxide (CO2)
Other components: molecular hydrogen (H2), methane (CH4), carbon monoxide (CO), ammonia (NH3), nitrous oxide (N2O), nitrogen dioxide (NO2), hydrogen sulfide (H2S), dimethyl sulfide [(CH3)2S], sulfur dioxide (SO2), and hydrogen chloride (HCl).
Some elements appear in multiple form—for example, carbon as carbon dioxide, methane, or dimethyl sulfide. It is useful to consider the occurrence of the elements before focusing on the more specific aspects of atmospheric chemistry (the forms in which the elements are present). One can speak of the Earth's “inventory of volatiles,” recognizing that the components of the inventory may be reorganized from time to time, but that it is always composed primarily of the compounds of hydrogen, carbon, nitrogen, and oxygen, along with the noble gases.

Processes
A process that delivers a gas to the atmosphere is termed a source for the gas. Depending on the question under consideration, it can make sense to speak in terms of either an ultimate source, the process that delivered a component of the volatile inventory to the Earth, or an immediate source, the process that sustains the abundance of a component of the present atmosphere. Any process that removes gas either chemically, as in the consumption of oxygen during the process of combustion, or physically, as in the loss of hydrogen to space at the top of the atmosphere, is called a sink.
Throughout the history of the atmosphere, sources and sinks have often been simultaneously present. While one process consumes a particular component, another produces it, and the concentration of that component in the atmosphere will rise or fall depending on the relative strengths of the sources and sinks. If those strengths are balanced (or nearly so), the composition of the atmosphere will not change (or will change only very slowly, perhaps imperceptibly); however, the molecules of the gas in question are passing through the atmosphere and are not permanently resident. The rate of the resulting turnover of molecules in the atmosphere is expressed in terms of the residence time, the average time spent by a molecule in the atmosphere after it leaves a source and before it encounters a sink.

Processes affecting the composition of the early atmosphere
Ultimate sources
The material from which the solar system formed is often described as a gas cloud or, at a later stage, solar nebula. The cloud was rich in volatiles (termed primordial gases) and must have been the ultimate source of the atoms in the present atmosphere. What is of primary concern, however, is the sequence of events and processes by which the volatiles present in the initial gas cloud were transferred to the Earth's inventory and the efficiency with which this was accomplished.
The formation of the solar system began when one portion of the gas cloud became dense enough due to compression by some external force—a shock wave from the explosion of a nearby supernova, perhaps—to attract gravitationally the material around it. This material “fell” into the region of higher density, making it even denser and attracting other material from still farther away. As gravitational collapse continued, the centre of the cloud became very dense and hot, because the kinetic energy of the incoming material was released as heat. Thermonuclear reactions began at the core of the central object, the Sun.

Capture and retention of primordial gases
Far from the central point, the material in the gas cloud tended to settle to an extensive equatorial plane around the Sun. As the material in this disk cooled, chunks of rock grew and accreted to form the planets. The planets are much less massive than the Sun, but if they grew large enough and if the gases around them were cool enough, they could accumulate an atmosphere from the volatile components of the gas cloud. A partial inventory of that cosmo-chemical stockpile, the starting point for atmospheric development, is shown in the column for the solar system in Table 1. This direct capture is the first of three source mechanisms that can be described.
A planetary atmosphere accumulated in this way would consist of primordial gases, but the relative abundances of the individual components would differ from those in the gas cloud if the gravitational field of the new planet were strong enough to hold some, but not all, of the gases around it. It is convenient to express the strength of a gravitational field in terms of escape velocity, the speed at which any particle (a molecule or spacecraft) must be traveling in order to overcome the force of gravity. For the Earth this velocity is 11.3 kilometres per second (seven miles per second), and it follows that, once the solid material of the Earth had accumulated, gas molecules passing the planet at lower speeds would have been captured and accumulated to form an atmosphere.
The speed at which a gas molecule moves is proportional to (T/M)1/2, where T is absolute temperature and M is molecular mass. As will be shown later, the uppermost layers of the present atmosphere are still very hot and might have been much hotter early in the Earth's history. At temperatures below 2,000 kelvins (K), however, molecules of any compound with a molecular weight greater than about 10 will have an average velocity of less than 11.3 kilometres per second. On this basis, it was long thought that the earliest atmosphere of the Earth must have been a mixture of the primordial gases with molecular weights greater than 10. Hydrogen and helium, with molecular weights of 2 and 4, should have been able to escape. Because hydrogen is the most abundant element in the solar system (see Table 1), it is thought that the most abundant forms of the other volatile elements were their compounds with hydrogen. If so, methane, ammonia, and water vapour, together with the noble gas neon, would have been the most abundant volatiles with molecular weights greater than 10 and, thus, the major constituents of the Earth's primordial atmosphere. The atmospheres of the four giant outer planets (Jupiter, Saturn, Uranus, and Neptune) are rich in such components, as well as in molecular hydrogen and, presumably, helium, which those more massive and colder bodies were apparently able to retain.

Outgassing of the solid planet
The release of gases during volcanic eruptions is one example of outgassing; releases at submarine hydrothermal vents are another. Although the gas in modern volcanic emanations commonly derives from rocks that have picked up volatiles at the Earth's surface and then have been buried to depths at which high temperatures remobilized the volatile material, a very different situation must have prevailed at the earliest stages of the Earth's history.
The planet accreted from solid particles that formed as the primordial gas cloud cooled. Long before the volatile components of the cloud began to condense to form massive solid phases (e.g., long before water vapour condensed to form ice), their molecules would have coated the surfaces of the solid particles of rocky material that were forming. As these solid particles continued to grow, a portion of the volatiles coating their surfaces would have been trapped and carried thereafter by the particles. If the solids were not remelted by impact as they collected to form the planet, the volatiles they carried would have been incorporated in the solid planet. In this way, even without collecting an enveloping gaseous atmosphere, a newly formed planet could include—as material occluded in its constituent grains—a substantial inventory of volatiles.
At some point in its early history, the Earth became so hot that much of the iron dispersed among the solid particles melted, became mobile, and collected to form the core. Related events led to the formation of rocky layers that were the precursors of the Earth's present-day mantle and crust. As part of this process of differentiation, volatiles present in the particles would have been released through outgassing. The outgassing must have occurred on a colossal scale if the accreting particles had retained their volatiles right up to the time of differentiation.
An atmosphere created by retention of these outgassing products would derive ultimately from nebular gases. Its chemical composition, however, would be expected to differ in two principal respects from that of an atmosphere formed by the capture of primordial gases: (1) whereas the captured atmosphere would contain all gases that were moving slowly enough (i.e., that were sufficiently cold and/or of sufficient molecular weight) so that it was possible for the planet to retain them gravitationally, the outgassed atmosphere would contain only those gases “sticky” enough to have been significantly retained in the rocky particles from which the planet formed; and (2) methane and ammonia, two presumed components of a captured atmosphere, would probably not be stable under the conditions involved in outgassing. Thus, the noble gases, which would be poorly held by particles, would be of low abundance relative to gases derived from chemically active elements. Further, the principal forms of carbon and nitrogen in an outgassed atmosphere would be carbon monoxide or carbon dioxide together with molecular nitrogen.

Importation
A compromise between the extremes of direct capture and outgassing proposes that the Earth's inventory of volatiles was delivered to the planet late in its accretionary history—possibly after differentiation was nearly complete—by impact of a “last-minute” crop of solid bodies that were very strongly enriched in volatile materials (these were the last substances to condense as the solar nebula cooled). Such bodies might have had compositions similar to those of comets that still can be observed in the solar system. These last-minute condensates may have coated the planet as a surface veneer that yielded gases only when heated during differentiation, or they may have released their volatiles on impact.
Because such bodies would have been relatively small, they would not have been able to retain primordial gases by means of a substantial gravitational field. Their complement of volatiles, retained by cold trapping in ices and on particle surfaces, would be expected to resemble the “sticky” (i.e., polar and reactive) gases occluded by solid particles at earlier stages of cooling of the gas cloud but possibly lost during earlier higher temperature phases of the Earth's accretion.

Sinks
The dominant pathways by which gases are removed from the present atmosphere are discussed below in the section Biogeochemical cycles. Apart from those processes, three other sinks merit attention and are described here.

Photochemical reactions
Sunlight can provide the energy required to drive chemical reactions that consume some gases. Due to a rapid and efficient photochemical consumption of CH4 and NH3, a methane–ammonia atmosphere, for example, would have a maximum lifetime of about 1,000,000 years. This finding is of interest because it has been suggested that life originated from mixtures of organic compounds synthesized by nonbiological reactions starting from methane and ammonia. Recognition of the short atmospheric lifetimes of these materials poses grave difficulties for such a theory. Water, too, is not stable against sunlight that has not been filtered by overlying layers containing ozone or molecular oxygen, which very strongly absorb much of the Sun's ultraviolet radiation. Water molecules that rise above these layers are degraded to yield, among other products, hydrogen atoms, H·.

Escape
Hydrogen and helium, or products like H·, tend to have velocities high enough so that they are not bound by the Earth's gravitational field and are lost to space from the top of the atmosphere. The importance of this process extends beyond the very earliest stages of the Earth's history because continuous sources exist for these light gases. Helium is continually lost as it is produced by the decay of radioactive elements in the crust.
A combination of photochemical reactions and the subsequent escape of products can serve as a source for molecular oxygen (O2), a major component of the modern atmosphere that, because of its reactivity, cannot possibly have derived from any of the other sources so far discussed. In this process, water vapour is broken up by ultraviolet light and the resulting hydrogen is lost from the top of the atmosphere, so that the products of the photochemical reaction cannot recombine. The residual oxygen-containing products then couple to form O2.

Solar-wind stripping
The Sun emits not only visible light but also a continuous flow of particles known as the solar wind. Most of these particles are electrically charged and interact only weakly with the atmosphere, because the Earth's magnetic field tends to steer them around the planet. Prior to the formation of the Earth's iron core and consequent development of the geomagnetic field, however, the solar wind must have struck the top layers of the atmosphere with full force. It is postulated that the solar wind was much more intense at that time than it is today and, further, that the young Sun emitted a powerful flux of extreme ultraviolet radiation. In such circumstances, much gas may have been carried away by a kind of atomic sandblasting that may have had a marked effect on the earliest phases of atmospheric development.

Biogeochemical cycles



Figure 1: A schematic representation of the biogeochemical cycle of carbon.
Interactions with the crust and, in particular, with living things, the biosphere, can strongly affect the composition of the atmosphere. These interactions, which form the most important sources and sinks for atmospheric constituents, are viewed in terms of biogeochemical cycles, the most prominent and central being that of carbon. The carbon cycle, outlined schematically in Figure 1, includes two major sets of processes: biologic and geologic.

Biologic carbon cycle
The biologic processes of photosynthesis and respiration mediate the exchange of carbon between the atmosphere or hydrosphere and the biosphere,
In these reactions, CH2O crudely represents organic material, the biomass of bacteria, plants, or animals; and A represents the “redox partner” for carbon (reduction + oxidation → redox), the element from which electrons are taken during the biosynthesis of organic material and which accepts electrons during respiratory processes. In the present global environment, oxygen is the most prominent redox partner for carbon (i.e., A = O in the above equation), but sulfur (S) also can serve as a redox partner, and modified cycles based on other partners (e.g., hydrogen) are possible. Imbalances in the biologic carbon cycle can change the composition of the atmosphere. For example, if O is the principal redox partner and if photosynthesis exceeds respiration, the amounts of O2 will increase. The carbon cycle can in this way serve as a source for O2. The strength of this source is dependent on the degree of imbalance between photosynthesis and respiration.
The biologic degradation of organic material and the release of products to the atmosphere need not involve an inorganic redox partner such as oxygen or sulfur. Communities of microorganisms found in sediments are capable of carrying out the process of fermentation, in which electrons are shuffled among organic compounds. Many individual steps catalyzed by a variety of organisms are involved, but the overall reaction amounts to
This process is an important source of atmospheric methane.

Geologic carbon cycle
The geologic portions of the carbon cycle can be described most conveniently by following a carbon atom from the moment of its injection into the atmosphere in the form of carbon dioxide released from a volcano. The carbon dioxide—any CO2 in the atmosphere—will come in contact with water in the environment and is likely to dissolve to form carbonic acid:
This weak acid is an important participant in weathering reactions that tend very slowly to dissolve rocks exposed to precipitation and groundwater at the Earth's surface. An exemplary reaction showing the conversion of a solid mineral to soluble products would be
where s indicates solid and aq stands for aqueous solution. Along with the other products of this reaction, the bicarbonate (HCO3-) derived from the volcanic CO2 would eventually be transported to the ocean. At all points in the hydrosphere, the bicarbonate would be in equilibrium with other forms of dissolved CO2 through chemical reactions that could be depicted as follows:
In settings where its concentration was enhanced, carbonate (CO32-) produced in this way could unite with the calcium ion Ca2+, which is naturally present in seawater due to weathering reactions, to form solid CaCO3, calcite, the principal mineral in limestone. The dissolved carbon dioxide might return to the atmosphere or remain in the hydrosphere. In either case, it eventually could enter the biologic carbon cycle and be transformed into organic matter. If the CaCO3 and the organic matter sank to the bottom of the ocean, they both would be incorporated in sediments and could eventually become part of the rocky material of the crust. Uplift and erosion, or very deep burial and melting with subsequent volcanic activity, would eventually return the carbon atoms of the CaCO3 and the organic matter to the atmosphere.

Interaction of biologic and geologic cycles
The pace of the biologic carbon cycle is measured in the lifetimes of organisms, while that of the geologic cycle is measured in the lifetimes of sedimentary rocks (which average about 600,000,000 years). Each interacts strongly with the atmosphere, the biologic cycle exchanging CO2 and redox partners and the geologic cycle supplying CO2 and removing carbonate minerals and organic matter—the eventual source of fossil fuels (e.g., coal, oil, and natural gas)—in sediments. An understanding of the budgets and pathways of these cycles in the present global environment enables investigators to estimate their effects in the past, when conditions (the extent of evolution of the biota, the composition of the atmosphere, and so on) may have been quite different.
The quantitative importance of these processes, now and over geologic time, can be summarized by referring to Table 2. Carbon in the atmosphere as carbon dioxide is almost the smallest reservoir considered in this tabulation, but it is the central point from which processes of the biogeochemical cycle have distributed carbon throughout the Earth's history. Reconstructions of atmospheric development must recognize that the very large quantities of carbon now found in sedimentary carbonates and organic carbon have flowed through the atmosphere and that the organic carbon (which includes all fossil fuels as well as far more abundant, ill-defined organic debris) represents material produced by photosynthesis but not recycled by respiration. The latter process must have been accompanied by the accumulation of the oxidized forms (e.g., molecular oxygen, O2) of carbon's redox partners.
Table 2 also emphasizes the dissolution of atmospheric gases by the ocean. The carbon dioxide in the atmosphere is in equilibrium with, and far less abundant than, carbon dioxide, bicarbonate ion (HCO3-), and carbonate ion (CO32- ) in the ocean. If all carbon dioxide were somehow suddenly removed from the atmosphere, the ocean would replenish the supply within a few thousand years (the so-called stirring time of the ocean). Likewise, any change in the concentration of CO2 in the atmosphere is accompanied by a quantitatively far larger change in the amount of CO2, HCO3-, and CO32- in the ocean. Similar equilibriums prevail for molecular nitrogen (N2) and molecular oxygen (O2). The atmosphere contains about 3,940,000 petagrams (Pg; one petagram equals 1015 grams) of nitrogen as N2, with about 22,000 Pg being dissolved in the ocean. Oxygen is distributed in such a way that 1,200,000 Pg of O2 are in the atmosphere while 12,390 Pg are in the ocean.

Weathering reactions
No matter what their origins, reactive gases in the atmosphere are likely to interact with other parts of the crust through what are termed weathering reactions. Not just carbonic acid associated with the carbon cycle but any acid becomes involved in acidic dissolution of susceptible rocks. As it does so, its concentration in the atmosphere declines, eventually reaching zero unless some process keeps replenishing the supply.
Even if respiration were suddenly to cease, oxygen produced by photosynthesis, or any oxidant in the atmosphere, would be consumed if oxidizable materials were present. The corrosion of metals is the most familiar example of this process in the modern world, but there are other examples involving natural forms of iron, sulfur, and carbon as well. Much of the iron bound in minerals is in the ferrous form (Fe2+). As this material is exposed by uplift and erosion, it consumes atmospheric oxidants to form ferric iron (Fe3+), the red, fully oxidized form of iron popularly identified as rust. Sulfide minerals (pyrite, or fool's gold, being the most familiar example) also consume oxidants as the sulfur is oxidized to produce sulfate. Finally, natural exposure of sedimentary organic matter, including coal beds or oil seeps, results in the consumption of atmospheric oxidants as the organic carbon is oxidized to produce carbon dioxide.

Sequence of events in the development of the atmosphere
Absence of a captured primordial atmosphere
If the planet grew large (and had, therefore, a substantial gravitational field) before all gases were dispersed from its orbit, it ought to have captured an atmosphere of nebular gases. The size and composition of such an atmosphere would depend on temperature as well as planetary mass. If the solid planet had reached full size and if temperatures were greater than 2,000 K, the minimum molecular weight that could be retained might have been high enough that the very abundant gases with molecular weights between 10 and 20 (methane, ammonia, water, and neon) would have been collected inefficiently, if at all. A thinner primordial atmosphere consisting of nebular gases with higher molecular weights (e.g., argon and krypton; see Table 1), however, ought still to have been captured.
In spite of this, characteristics of the present atmosphere (see below) show clearly that a primordial atmosphere either never existed or was completely lost. Explanations offered for both of these possibilities are linked to the development of the Sun itself. Astronomical observations of developing stars (i.e., bodies similar to the early Sun) have shown that their early histories are marked by phases during which the gas in their surrounding nebulas is literally blown away by the pressure of light and particles ejected from the stars as they “turn on.” (After this initial intense activity, young stars begin life with an energy output significantly below their mid-life maximum.) If the removal of gases occurred in the solar system after involatile solids had condensed but before the inner planets (Mercury, Venus, Earth, and Mars) accreted, it would have been impossible for the Earth to capture a primordial atmosphere. Alternatively, if planetary accretion preceded ejection of gases and the Earth had accumulated a primordial atmosphere, perhaps the early solar radiation, particularly the solar wind, was so intense that it was able to strip all gases from the inner planets, meeting the second condition described above—namely, complete loss.

Secondary atmosphere
The atmosphere that developed after primordial gases had been lost or had failed to accumulate is termed secondary. Although the chemical composition of the atmosphere has changed significantly in the billions of years since its origin, the inventory of volatile elements on which it is based has not.

Origin
The elemental composition of the volatile inventory reveals its secondary origin. Abundances are given in Table 1 for 12 nuclides that can be associated with four groups: (1) chemically active volatiles (H, C, N, O, S), (2) primordial noble gases (4He, 20Ne, 36Ar, 84Kr), (3) elements that form involatile minerals (O, Mg, S, Fe), and (4) a noble gas derived by the radioactive decay of an involatile element (40Ar, derived from potassium). A comparison of entries in Table 1 shows that these groups have been collected by the Earth with sharply varying efficiencies. The column headed “collection efficiency” has been derived by the division of the abundance of each element on Earth by its abundance in the solar system and multiplying by 100. If the collection efficiency is close to 100 percent, the abundances are nearly equal and the transfer of this element from the solar system's initial reservoir to the planet was highly efficient. If the collection efficiency is low, most of the element was lost and is “missing” from the Earth's inventory. It is evident from Table 1 that efficiencies of collection are correlated primarily with chemical characteristics, not mass. This is the pattern expected if volatiles were retained by chemical interactions that yielded involatile phases rather than by gravitational attraction. Collection efficiencies for O, Mg, S, and Fe (which are included here only as representatives of the broad range of elements that were largely bound in involatile solid phases as the solar nebula cooled) are high. Those for the chemically active volatiles that could not form minerals stable at high temperatures (H, C, and N) are much lower. Spectacularly decreased efficiencies of collection are associated with the primordial noble gases.
The evidence points decisively to a process in which the elements to be retained in the terrestrial inventory were separated from those to be lost by a separation of solids from gases. The chemically active volatile elements could be incorporated in solids by formation of nitrides and carbides, by hydration of minerals, and by inclusion in crystal structures (e.g., as ammonium and hydroxide ions) and could form some relatively involatile materials independently (organic compounds with high molecular weights are found in meteorites and were probably abundant in the cooling solar nebula); yet, none of these mechanisms was available to the noble gases. Formation of a group of solids rich in chemically active volatiles but not large enough to retain noble gases, followed by a loss of all materials still in the gas phase and an incorporation of the volatile-rich solids in the Earth, would be consistent with the chemical evidence and with the processes described above as outgassing and importation.
The special case of 40Ar (Table 1) is particularly indicative of the derivation of the atmosphere through outgassing. Whereas the other noble-gas isotopes noted in Table 1 (4He, 20Ne, 36Ar, 84Kr) are primordial in origin, 40Ar derives primarily from the radioactive decay of the isotope potassium-40. Therefore, even though the solar system abundance of 40Ar is much lower than that of 36Ar, its abundance on Earth is much higher because, uniquely among the noble-gas isotopes listed in Table 1, its source—the rock-forming element potassium (K)—is part of the solid planet. As radioactive potassium in rocks decayed over the Earth's history, the 40Ar produced first became trapped within mineral crystals at sites formerly occupied by K+, then was released when the crystals were melted in the course of igneous activity, and eventually reached the surface through outgassing. Given the abundance of potassium in the Earth's crust, it would be impossible to attribute the origin of the atmosphere to outgassing if the abundance of 40Ar was far lower than that of 36Ar, as in the solar system.

Early composition
The most critical parameter pertaining to the chemical composition of an atmosphere is its level of oxidation or reduction. At one end of the scale, an atmosphere rich in O2 (like the present one of the Earth) is termed highly oxidizing, while one containing molecular hydrogen, H2, is termed reducing. These gases themselves need not be present. Modern volcanic gases are located, for example, toward the oxidized end of the scale. They contain no O2, but all hydrogen, carbon, and sulfur are present as H2O, CO2, and SO2 (oxidized forms), while nitrogen is present as N2 (not NH3). A relationship prevails between the oxidation or reduction of outgassing volatiles and the inorganic material with which they come in contact: any hydrogen, carbon, or sulfur brought into contact with modern crustal rocks at volcanic temperatures will be oxidized by that contact.
The abundance of hydrogen in the solar nebula, the common occurrence of metallic iron in meteorites (representative of primitive solids), and other lines of geochemical evidence all suggest that the Earth's early crust was much less oxidized than its modern counterpart. Although all iron in the modern crust is at least partly oxidized (to Fe2+ or Fe3+), metallic iron may have been present in the crust as outgassing began. If the earliest outgassing products were equilibrated with metallic iron, hydrogen would have been released as a mixture of molecular hydrogen and water vapour, carbon as carbon monoxide, and sulfur as hydrogen sulfide. The presence of metallic iron during the last stages of outgassing is, however, unlikely, and, because H2 is not gravitationally bound, it would have been lost rapidly. At an early point, hydrogen would have been almost completely in the form of water vapour and carbon in the form of carbon dioxide. Nitrogen would have been outgassed along with the carbon and hydrogen. As carbon dioxide was consumed by weathering reactions and water vapour condensed to form the oceans, molecular nitrogen must have become the most abundant gas in the atmosphere. It is certain that molecular oxygen was not among the products of outgassing.
Among the oldest rocks are water-laid sediments with an age of 3,800,000,000 years. Neither they nor any other ancient rocks contain metallic iron, though nearly all contain oxidized iron (Fe2+). Carbon is present both as organic material and in a variety of carbonate minerals. The existence of these sediments requires atmospheric pressures and temperatures consistent with the presence of liquid water. The nature of the iron minerals and their abundance suggest that Fe2+ was a significant component of ocean water and that concentrations of O2 had to have been essentially zero because Fe2+ reacts very rapidly with O2.
The presence of organic carbon and carbonate minerals in the sediments dated 3,800,000,000 years old would be consistent with the development of a biologically mediated carbon cycle by that point in time, but the degree of preservation of these materials (which were heated to temperatures near 500° C [932° F] for millions of years at some point in their history) is so poor that the question cannot be settled. Relatively well-preserved sediments with an age of 3,500,000,000 years are far more abundant. In addition to abundant organic carbon and carbonate minerals, they contain microfossils and sedimentary features demonstrating convincingly that life had arisen on Earth by that time. The distribution of the stable isotopes of carbon (carbon-12 and carbon-13) in sedimentary materials younger than 3,500,000,000 years demonstrates that living organisms were effectively in control of the global carbon cycle from that time onward.
The existence of sedimentary carbonates is direct evidence that carbon dioxide was present in the atmosphere. Its precise abundance is not known, but the best estimates are that it was substantially, perhaps 100 times, higher than the present atmospheric level. A strongly enhanced greenhouse effect (see below Atmospheric heat budget and energy transfer), leading to more efficient retention of heat derived from solar radiation, would be expected. For many students of the Earth's history, the fact that the early oceans did not freeze in spite of the dim Sun is evidence that the abundance of atmospheric carbon dioxide was high enough to provide the enhanced greenhouse effect.

Rise of molecular oxygen



Figure 2: A “best guess” reconstruction of the abundance of O2 in the …
Recognition of the nature of the Earth's pre-oxygenic environment is critical to consideration of this problem. If humans could somehow take a spaceflight not to another planet but to the Earth of 3,000,000,000 years ago, they would find that space suits would have been required on their home planet at that time. More dramatically, if those time-traveling astronauts were somehow able to take with them all of the oxygen from the modern atmosphere, they would find that it would disappear soon after release. Not only was oxygen absent in the early atmosphere but potent sinks for O2 were abundant as well. Oxidizable materials such as ferrous iron, sulfides, and organic compounds littered environments in which they are now absent. These chemicals absorbed O2 almost immediately after its release. Moreover, as the oxygen-absorbing capacity of such compounds was exhausted, new material that had been eroded from the unoxidized crust took their place. This process continued until the rock cycle (sedimentation, burial, igneous activity, uplift, and erosion) had exposed all oxidizable materials in the crust. No matter what the supply of O2, the process must have taken time (about half the rock volume of the crust is recycled every 600,000,000 years). It is, therefore, very important to distinguish clearly between the first biologic production of O2 and its persistent accumulation in the atmosphere. It is conceivable, even likely, that these events were separated by hundreds of millions of years. Evidence bearing on the oxygenation of the Earth's atmosphere is summarized in Figure 2 and discussed in the following sections. The abundance of O2 at each point is expressed in terms of its approach to the present atmospheric level (PAL). For example, because the pressure of O2 in the present atmosphere is 0.21 atm (abbreviation for atmosphere, a unit of pressure equal roughly to 14.7 pounds per square inch), a planetary atmosphere containing 10 percent of that amount, 0.021 atm, would be described as having an oxygen level of 0.1 PAL.

Photochemical production
The strength of this source is limited by the requirement that water vapour rise in the atmosphere to altitudes at which solar ultraviolet radiation capable of cleaving water molecules has not yet been absorbed by other atmospheric constituents. The transport of water vapour to high altitudes is severely impeded by a cold layer in the atmosphere. Water vapour freezes in this layer, and the rate of photochemical production of O2 is thus limited. The severity of this limitation is not precisely known, but it is evident that atmospheric levels of oxygen did not rise until oxygenic photosynthesis was well established. This does not indicate that photochemical production of O2 was insignificant. Rather, it demonstrates that the strength of the process as a source was exceeded by the strength of the contemporary oxygen sinks (chiefly oxidative weathering reactions at the Earth's surface) and that residence times for O2 were so short that significant atmospheric concentrations could not accumulate. The best estimate is that pressures of O2 at sea level and ground level were less than 5 × 10-8 PAL.

Onset of oxygenic photosynthesis
The development of a biologically mediated carbon cycle prior to 3,500,000,000 years ago virtually requires that some form of photosynthesis had arisen by that time, but the possibility remains that sulfur or hydrogen, not oxygen, was serving as the redox partner. It also has been noted that some sediments 3,500,000,000 years in age contain microfossils with shapes resembling those of modern oxygenic photosynthesizers. This is suggestive, though not compelling, evidence that oxygenic photosynthesis had developed by 3,500,000,000 years ago. Shape is an infamously imprecise indicator of biochemical characteristics of microorganisms. More specifically, while it might be possible to recognize a photosynthetic organism from its shape, it is very difficult to determine exactly what redox partners that organism employed.
Geochemical and paleontological features of sedimentary rocks 2,800,000,000 years in age offer stronger evidence that oxygenic photosynthesis had arisen by that time. At 2,800,000,000 years, the abundance of carbon-13 in sedimentary organic carbon decreases sharply from levels maintained between 3,500,000,000 and 2,900,000,000 years ago, then slowly rises, regaining those levels about 2,200,000,000 years ago. This has been interpreted in terms of a transient in the biogeochemical carbon cycle in which biogenic methane (which is strongly depleted in carbon-13) served as an important mobile constituent of the cycle during the interval from 2,800,000,000 to 2,200,000,000 years ago. According to this interpretation, methane was able to play this role only after O2 became available and facilitated its metabolism. As O2 sinks decreased in strength and the atmosphere became oxidizing, however, the mobility of methane was reduced and the methane cycle took on its modern form, which seldom leads to strongly decreased abundances of carbon-13 in sedimentary organic matter.
Microfossils resembling modern oxygenic photosynthesizers also appear in sediments of this age, and they are accompanied by sedimentologic features (apparent “fossil gas pockets”) that are interpreted as evidence of aerobic metabolism. Thus, evidence dating from about 2,800,000,000 years ago is more abundant and diverse (geochemically, morphologically, and sedimentologically) than that found in rocks 3,500,000,000 years of age. In spite of these points of consistency, this evidence is not decisive.
Evidence from younger sediments (see below) indicates that oxygenic photosynthesis almost certainly developed earlier than 2,200,000,000 years ago. Whatever the precise moment of development, it marked the origin of the first so-called oxygen oasis, a restricted environment in which the abundance of O2 rose above 5 × 10-8 PAL probably quite significantly. Within such oases, aerobic metabolism could occur. At their margins, the delivery of oxidizable materials from the surrounding global environment overwhelmed the local supply of O2. Overall, the atmosphere did not become oxidizing, but, as oxygenic photosynthesizers proliferated, the number and size of the oases grew.

Transition to an aerobic environment
Pyrite and uraninite are minerals of iron and uranium, respectively, that are not stable in the presence of O2. Though they can be found in some modern river sediments, neither can survive in them for thousands of years. Yet, many sediments older than about 2,200,000,000 years contain well-rounded grains of these minerals. Their shapes and locations indicate prolonged exposure and tumbling in ancient rivers or as beach deposits, but there is no evidence of chemical attack by oxygen. The precise significance of this observation is best considered together with measurements of the movement of iron in fossil soil profiles.
If soil gases (in equilibrium with the atmosphere) contain O2, iron exposed during the breakdown of soil minerals will be immobilized by oxidation and will not be leached from near-surface soil horizons. Conversely, if O2 is absent during soil development, chemical analysis of fossil soils will reveal depletion of iron near the former soil surface. Rates of the dissolution of uraninite and the leaching of iron in soil profiles also depend on the abundance of carbon dioxide (CO2). Because the patterns of dependence are different, the combination of evidence based on both phenomena allows for the estimation of abundances of both CO2 and O2. This line of interpretation leads to the conclusion that about 2,200,000,000 years ago the ratio of the molecular abundance of O2 to that of CO2 was about 1.3 (at present it is 635), and that the pressure of O2 was near 0.01 PAL while that of CO2 was about nine times higher than at present. Other workers agree that the uraninite and fossil soil data indicate the development of oxidizing conditions at the surface by 2,200,000,000 years ago, but they place the most probable level of O2 lower by a factor of 10 or more.
The consumption of oxidizing power by the crust is recorded by the inorganic constituents of sedimentary rocks. Iron-bearing sediments, or iron formations, are of particular interest because the collection of substantial quantities of iron in a sedimentary basin requires that iron be mobile in the world ocean. Mobility requires solubility, and, while Fe2+ is soluble, Fe3+, the form of iron that results if O2 comes in contact with Fe2+, is highly insoluble.
Three states can be distinguished: (1) The existence of iron formations containing only Fe2+ suggests a complete absence of oxygen. (2) The existence of iron formations containing Fe2+ and Fe3+ indicates that levels of oxygen were low enough—essentially zero in the deep ocean—so that iron was mobile, but it also suggests that O2 (perhaps at an oxygen oasis) was important in triggering deposition of the iron, though other means of oxidation—photochemical processes, for example—are quite conceivable. (3) The disappearance of iron formations from the sedimentary record suggests persistent oxygenation of the ocean. This sequence of possibilities is represented in the geologic record as follows: (1) The oldest sedimentary rocks are iron formations that contained only Fe2+ at the time of their deposition. (2) The first appearance of primary Fe3+ (produced during the formation of the rock rather than in later weathering) was in iron formations about 2,700,000,000 years ago. (3) Iron formations disappeared almost completely from the record about 1,700,000,000 years ago (with a few isolated and very small recurrences about 1,000,000,000 years ago). Moreover, the abundance of iron formations increased significantly from 2,700,000,000 to 2,200,000,000 years ago, suggesting that some new factor, possibly oxidative precipitation of Fe3+, was enhancing the rate of deposition. It is for this same time interval that carbon-isotopic evidence indicates the operation of an O2-dependent methane cycle.
Evidence for the evolution of eukaryotic organisms (those containing a membrane-bound nucleus and other organelles) first appears in the microfossil record of about 1,400,000,000 years ago. Biochemical reactions that occur during the growth and division of such cells require oxygen levels of 0.02 PAL. Attainment of that level by 1,400,000,000 years ago apparently led to oxygenation of the deep sea and the cutoff of deposition of iron formations about 1,700,000,000 years ago.

Attainment of the modern O2 level
The abundance of carbon-13 in sedimentary organic materials and in carbonates from 900,000,000 to 600,000,000 years ago indicates that unusually large quantities of organic carbon were buried without reoxidation during that interval. The burial of this carbon must have been accompanied by the accumulation of oxidized forms of carbon's redox partners. The quantities released were adequate to raise the level of O2 to 1.0 PAL or more.
It has been calculated that oxygen requirements of the earliest animals, which developed about 700,000,000 years ago, would have been met—if the animals had circulatory systems that incorporated oxygen carriers like hemoglobin—by O2 abundances as low as 0.1 PAL. If circulatory systems had not yet evolved, an O2 abundance of 1.0 PAL would have been required. Studies of fossils indicate that the animals were very thin (one to six millimetres [0.04–0.24 inch]) in spite of great breadth and length (up to 1,000 millimetres [39 inches]). Such a shape seems optimized for transport of O2 by diffusion from the surrounding water to the cells in which it was needed, thus pointing to the latter higher value (namely, an O2 abundance of 1.0 PAL). Other reconstructions of O2 levels based on biologic evidence suggest that the widespread development of land plants about 400,000,000 years ago must have driven O2 to levels near 1.0 PAL, and they show O2 levels rising smoothly from levels near 0.1 PAL at 650,000,000 years ago to 1.0 PAL at 400,000,000 years ago.

Variation in abundance of carbon dioxide
The approximately hundredfold decline of atmospheric CO2 abundances from 3,500,000,000 years ago to the present has apparently not been monotonous. During that interval, numerous ice ages have come and gone. Significant changes in climate can result from geographic changes, but it is generally concluded that modulation of the efficiency of the Earth's greenhouse effect is also required to produce the extreme variations associated with widespread continental glaciations. In recognition of this, broad climatic variations during the past 750,000,000 years have been described in terms of alternating “icehouse” and “greenhouse” episodes.
Icehouse conditions—apparently associated with the depletion of atmospheric CO2, the principal greenhouse gas—have prevailed since about 65,000,000 years ago and during two earlier periods, 650,000,000–530,000,000 and 360,000,000–240,000,000 years ago. It is suggested that intervening greenhouse episodes have been associated with higher abundances of CO2 in the atmosphere. The hypothesis is far from proved. Nonetheless, its details are being explored aggressively amid concerns that the accelerated production of CO2 due to industrial combustion of fossil fuels may reverse climatic conditions with catastrophic effect. It is feared that such a reversal could result in the melting of the polar ice caps and a consequent flooding of coastal areas (see below Climate modification).

J.M. Hayes
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 wavelengths—namely, 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.

Atmospheric heat budget and energy transfer



Figure 5: Energy budget for the surface of the Earth illustrating what happens, on average, to 100 …
The overall heat budget of the atmosphere and surface is summarized in Figure 5. The surface, on average, receives 17 percent of its heat directly from the Sun, 15 percent from solar radiation scattered by clouds, and the balance, 68 percent, from absorption of infrared radiation emitted by the atmosphere. The greater part of the energy absorbed by the surface, 79 percent, is returned to the atmosphere in the form of radiation, with spectral properties determined by the local ground temperature. The remainder, 21 percent, is transmitted to the atmosphere by conduction and as a by-product of the exchange of water, H2O. The surface can cool by evaporation of H2O, and the associated heat is transmitted to the air as vapour, which recondenses to form clouds and either rain or snow. Phase changes of H2O play a major role in the energy budget of the lower atmosphere. It is, in fact, the importance of H2O that sets the Earth apart from all of its neighbours in the solar system.
The atmosphere can be conceived of as a compressible fluid of infinite extent that is heated from below by a moist radiating surface and perturbed locally by energy absorbed from sunlight. Direct absorption of solar radiation is important primarily for the stratosphere and thermosphere. Transfer of energy by infrared radiation is a dominant mode for heat transmission between the different atmospheric layers, with an additional contribution due to motions generated by spatial heterogeneities in heating rates.

Greenhouse effect
Transfer of energy by radiation is effected mainly by trace constituents of the atmosphere, primarily H2O, CO2, and O3. In contrast to the major constituents, N2 and O2, these gases are able to absorb the longer wavelengths of the planetary radiation field. They thus assume an importance out of proportion to their abundance. They act to trap heat radiated by the surface, much as the glass panes of a greenhouse do. Like the glass, the atmosphere is transparent to sunlight but is essentially opaque to longer wavelengths. The infrared-active gases return heat to the ground, accounting for about 70 percent of the net input of energy to the surface. If the atmosphere were devoid of water and carbon dioxide, the surface temperature would be about 40 K colder than it is today, and large portions of the planet would be covered by ice.
Since the early 1980s there has been growing concern over the possibility that an increase in the abundance of carbon dioxide caused by combustion of fossil fuels could lead to a general warming of the global climate. Similar effects can arise from increases in the abundances of methane, nitrous oxide, and various chlorofluorocarbons (CFCs) such as CCl2F2 and CCl3F. These species are referred to collectively as the greenhouse gases in recognition of their ability to trap heat. Their importance to climate is explored in greater detail below in Climate modification.

The role of ozone
The increase of temperature with altitude above the tropopause is due primarily to absorption of solar radiation by ozone, O3. As noted above, O3 is a minor constituent of the atmosphere, a product of the interaction of molecular oxygen, O2, with sunlight. It plays a critical role in the global life-support system, however, absorbing most of the light incident on the Earth in the ultraviolet portion of the spectrum with wavelengths between about 200 and 300 nanometres (one nanometre [nm] equals 10-9 metre). The absorption process results in the dissociation of O3. A portion of the photon energy (hν) is deposited directly as heat; a larger fraction appears initially as internal excitation of O and O2. The internal energy is degraded rapidly to heat in the stratosphere by collisions with O2 and N2. Finally, the chemical potential energy represented by O + O2 is itself converted to heat as O3 is reformed by reaction of O with O2.
The recombination process involves two separate reactions,
where O3* denotes an unstable, energetic intermediate form of O3. The excess energy in O3* is removed by collisions with atmospheric molecules M in the second step. The pair of reactions is usually considered as a single reaction,
The rate of recombination, the number of O3 molecules formed per unit volume per unit time, is proportional to the product of the concentrations of O, O2, and M. The constant of proportionality, the reaction rate constant, may be determined experimentally in the laboratory.
To an excellent approximation in the stratosphere, it may be assumed that all of the radiative energy absorbed by O3 is converted locally to heat. The heating rate—and ultimately the course of temperature with altitude—depends on the details of the distribution of O3 with height.

Temperature inversion in the thermosphere
The inversion of temperature above 80 kilometres, in the thermosphere, is due to energy extracted from sunlight at wavelengths below 200 nanometres. There are several important processes. Between 100 and 200 nanometres, absorption of solar radiation leads to dissociation of O2,
At shorter wavelengths, absorption is associated primarily with the ionization of O, O2, and N2,
(Here, hν represents a photon and e is the electron emitted.) Integrated over altitude, dissociation of O2 is balanced by recombination. Molecular oxygen is reformed, either directly by
or indirectly by reactions such as
Sequence (6) is equivalent in effect to reaction (5). It can proceed more rapidly, however, under appropriate conditions in the atmosphere, depending on the efficiency of the step associated with the production of O3 and the availability of hydrogen, H. The recombination path (5) is said to be catalyzed by the presence of H. As will be seen, catalytic reaction chains play a major role in much of the chemistry of the atmosphere below 80 kilometres.
The path from the absorption of solar energy by reactions (3) and (4) to the ultimate disposal of this energy as heat is less direct for the thermosphere than for the stratosphere. Recombination of O atoms, mainly by reactions (5) and (6), can occur at altitudes quite different from those at which O2 is dissociated. The lower densities of the thermosphere allow O atoms to diffuse downward, with a compensating upward flow of O2. Recombination requires relatively high densities and is confined mainly to altitudes below 100 kilometres. Dissociation of O2, on the other hand, can proceed at any level, limited solely by the supply of O2 and by the availability of photons with energy sufficient to fragment the bond in O2. This spatial separation of dissociation and recombination results in a vertical redistribution of energy. Energy absorbed at one level by O2 is converted to heat at another and is transferred in between in the form of chemical potential energy represented by a downward flow of O. Further, only a fraction of the available chemical potential energy is converted to heat. The balance is released as radiation, contributing to the phenomenon of airglow (see below).
The fate of the energy deposited in (4) is similarly complex. A fraction of the photon energy is imparted initially to the photoelectron (an electron ejected by the action of a single photon). The photoelectron loses energy by collisions, both elastic and inelastic, with atmospheric molecules. It can cause further ionization and contribute to the production of excited states and the associated emission of airglow. Photoelectrons share energy efficiently with ambient electrons. As a consequence, the electron temperature tends to rise above the temperature of the neutral gas. There is a flow of energy from electrons to ions to neutrals. To some extent the thermosphere must be treated as though it were composed of three characteristic thermal reservoirs, each one having its own distinct temperature.
The heating rate depends on a complex suite of atomic and molecular processes and is difficult to evaluate precisely. To further complicate matters, the chemical potential energy represented by the photo-ion can contribute to heating at levels far removed from its original source.
Electrons are removed by dissociative recombination of molecular ions such as the oxygen ion O2+ and that of nitric oxide, NO+:
A portion of the potential energy is converted to heat in reactions (7) and (8). Some is used to produce excited, radiating states of O and N, and some is stored in excited states and degraded to heat subsequently by collisions. The balance is represented by the atoms N and O and is released only when the atoms recombine at a lower altitude to reform the stable molecules N2 and O2.
An atom or molecule of mass m is bound to the Earth by the force of gravity. The work done in moving a vertical distance ΔZ is mg ΔZ, where g defines the gravitational acceleration, 9.8 metres per second per second. The work that must be done to escape the gravitational field is mgR, where R is the radius of the Earth—6,400 kilometres. The average kinetic energy of the atom is (3/2)kT, where k is the Boltzmann constant, 1.38 × 10-16 erg per degree Kelvin and T is temperature. The atom can escape the gravitational field if kT is much larger than mgR. In practice, escape is impeded by collisions, except at the highest levels where the density of the atmosphere is low and collisions are comparatively rare.

Scale height
The atmosphere extends to great heights, with density declining by a factor of e (2.72) over an altitude interval given by (kT)/(mg). The quantity (kT)/(mg) is known as the scale height, denoted customarily by H. As noted above, it is proportional to the ratio of thermal kinetic energy to gravitational potential energy and measures the capacity of the atmosphere to support itself thermally in opposition to the confining force of gravity. The scale height is about seven kilometres in the lower atmosphere, rising to values in excess of 50 kilometres in the thermosphere where the mean molecular mass is smaller, reflecting the importance of lighter gases such as oxygen and helium. The scale height provides a useful measure of the vertical extent of the atmosphere: the atmosphere above a height z contains a total mass equivalent to that represented by a hypothetical layer of vertical extent H containing gas with density equal to the density at z.
If the mean distance traveled by an atom or molecule between collisions exceeds H, it can be assumed that the effect of collisions will be relatively unimportant. The average distance between collisions, known as the mean free path and denoted by λ, is given by {Qn}-1, where Q is the collision cross section, the target area offered by a typical atom or molecule, and n is the number of atoms or molecules (targets) per unit volume. The level at which λ = H is defined as the critical level, zc. Effects of collisions may be ignored for heights greater than zc. The region above zc is known as the exosphere. An atom or molecule moving upward through zc with speed greater than about 11 kilometres per second has kinetic energy in excess of its gravitational potential energy and in the absence of collisions is free to escape from the Earth.

Hydrogen loss
Atoms or molecules of a particular species have a range of speeds determined by their temperature, as described by the Maxwell–Boltzmann distribution law. For temperatures near the critical level, 750–2,000 K, significant numbers of hydrogen atoms have velocities greater than 11 kilometres per second. These atoms readily escape. There is an upward flow of hydrogen at all levels of the atmosphere to supply the flux leaving the Earth. Hydrogen is lost at a rate of about 108 atoms per square centimetre per second averaged over the surface of the planet. These escaping hydrogen atoms are derived ultimately from H2O in the oceans. Integrated over the 4.5 × 109 years of the Earth's history, the escape of hydrogen has removed about two metres depth of H2O from the surface of the oceans, liberating a quantity of O2 roughly equivalent to that present in the atmosphere today. The escape of hydrogen is not, however, thought to be the major factor in regulating the level of atmospheric O2. Burial of organic carbon in sedimentary rocks provides a more immediate source, though the escape of hydrogen may have been important at an earlier stage in the life of the Earth.
Thermal escape, as discussed here, is important mainly for hydrogen. There is, however, also a significant loss of helium. The loss mechanism in this case is believed to involve positively charged helium ions, He+, formed by photoionization. These ions are thought to be accelerated by electric fields and ejected from the Earth at higher latitudes where magnetic field lines are, on occasion at least, open to space. Helium is a transitory constituent of the atmosphere. It originates in the Earth's crust and is quickly lost to space. Its concentration reflects a dynamic balance between the source from within and loss to the outside.

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 gases—e.g., N2 and O2—are distributed more or less homogeneously around the Earth. The shorter lived species—e.g., carbon monoxide, nitric oxide, and ozone—can 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 productivity—less 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.

Nitrogen compounds
In contrast to carbon, which finds its most stable home in sediments, most of the Earth's nitrogen is in the atmosphere as the relatively inert gas N2. The N–N bond in N2 is very strong and is not easily fractured in the atmosphere, except at high altitudes where the molecule is exposed to energetic ultraviolet radiation or at low altitudes where it may be raised to a temperature exceeding 2,000 K near a lightning stroke. The bond can be broken by biologically mediated reactions, however, by photosynthetic blue-green algae or by bacteria functioning in symbiosis with plants of the legume family. Dissociation of N2 is essential for life: amino groups, represented by NH2, are indispensable components of living tissue. The N–N bond must be broken before the atom can be incorporated in living organisms. The class of compounds containing odd numbers of N atoms is referred to collectively as fixed nitrogen. When the bond in N2 is broken, the molecule is said to be fixed.

Fixed nitrogen



Figure 8: Major paths for oxidation and reduction of nitrogen. The influence of terrestrial life on …
Fixed nitrogen occurs in a variety of oxidation states, as shown in Figure 8. Its importance for the biosphere may be attributed in no small measure to this aspect of its chemistry. Oxidation of ammonium, NH4+, represented by
involves a change in free energy of 51.8 kilocalories at 298 K and one atmosphere. It denotes the first step in the oxidation of nitrogen, primary nitrification, in which nitrogen is transformed from oxidation state -3 to +3. The energy liberated in reaction (11) is utilized by bacteria, Nitrosomonas and similar genera, which effect the reaction. A different group of bacteria, Nitrobacter, is able to oxidize nitrogen further, from nitrite (NO2-) to nitrate (NO3-),
a process known as secondary nitrification. Reaction (12) involves a release of free energy of 20.1 kilocalories per mole.
Plants can use the ammonium ion, NH4+, or NO2- or NO3- to satisfy their need for fixed nitrogen but generally tend to prefer NH4+. Nitrogen in NO2- and NO3- must be reduced from oxidation state +3 or +5 to -3 before it can be assimilated into living tissue. The reactions—in this case assimilatory reduction—are denoted by
and
The necessary source of energy here is supplied by oxidation of organic matter, represented by CH2O.
The atmosphere contains 4 × 1015 metric tons of nitrogen in the form of N2. Approximately 108 metric tons are fixed annually by biologic agents, and about the same amount is transformed by various industrial processes associated with the combustion of fossil fuels and the manufacture of chemical fertilizer. In the absence of a source, the atmosphere would lose its reservoir of N2 in about 2 × 107 years. The gas would be oxidized to NO3- and would tend to accumulate in the oceans. This obviously has not happened. There is a return of nitrogen from NO3- to N2 through bacterially mediated denitrification, which is described by the reaction
Denitrification proceeds under anaerobic conditions (i.e., in the absence of free oxygen) and provides another example of the splendid efficiency of the global life-support system. Waste for one species is opportunity for another. Reaction (15) is exothermic: it releases energy equivalent to 124 kilocalories per mole. It represents an essential link in the nitrogen cycle, providing a respiratory path for bacteria when O2 is deficient and serving at the same time to maintain a relatively constant level of atmospheric N2.




Nitrous oxide, N2O, is the second most abundant nitrogen constituent of the atmosphere, as indicated in Table 3. It is formed as a by-product of denitrification, by reduction of NO3- and by oxidation of NH4+, the first step in nitrification. Additional production of N2O occurs during fossil fuel combustion. The lifetime of N2O in the atmosphere is about 150 years; it is removed mainly by photolysis (chemical decomposition by the action of radiant energy) in the stratosphere. Given its relatively short lifetime, as compared with N2, one might expect N2O to vary, and indeed there is evidence for a contemporary increase in the concentration of the gas in the atmosphere. This apparent increase is attributed in part to the burning of fossil fuels and in part to the general anthropogenic disturbance of the global nitrogen cycle. The human role in nitrogen fixation is now comparable to that of nature. Other links in the nitrogen cycle might be expected to adjust accordingly. The atmosphere contains various forms of fixed nitrogen besides N2 and N2O. As mentioned above, fixed nitrogen is introduced into the air by lightning at a rate of about 107 tons N yr-1 (metric tons of nitrogen per year). It also is released as a product of the combustion of fossil fuels and the burning of biomass. In addition, fixed nitrogen emanates from soil as a result of the microbial oxidation of NH4+, and it is formed in the stratosphere by a process involving the reaction of the metastable oxygen atom O(1D) with N2O,
In all cases, the primary source of fixed nitrogen is nitric oxide, NO. About 5 × 107 tons N yr-1 of NO are introduced directly into the troposphere, much of this in continental areas, with a particularly large contribution from industrial sources and automobiles and trucks in cities. The stratospheric source is smaller, about 106 tons N yr-1.
Reactions in the atmosphere result in a variety of secondary nitrogen species. Important compounds in remote regions include nitrogen dioxide (NO2), the unstable intermediate nitrogen trioxide (NO3), dinitrogen pentoxide (N2O5), nitrous acid (HNO2), and nitric acid (HNO3). Significant, too, is the set of reactive nitrogen compounds, NO + NO2 + NO3 + N2O5 + HNO2 + HNO3 + NO2NO2, designated NOx. The subset NO + NO2 plays a particularly important role in the photochemistry of ozone and is identified separately as NOy.

PAN and its role
The compound CH3CO.O2NO2, known as peroxyacetyl nitrate, or PAN, is thought to provide an important storage reservoir for NOx in the unpolluted troposphere. It decomposes thermally in the troposphere, with production of NO2 and the peroxyacetyl radical, CH3CO.O2,

Special Comp





Reaction (17) describes an equilibrium, which favours PAN at low temperatures and products at high temperatures. Conditions are such that PAN is relatively stable in the upper troposphere, but it decomposes rapidly in the warm environment near the ground. The compound is believed to form in urban air, a product of the complex set of reactions involved in the oxidation of hydrocarbons. From there, it can be exported via air motions to the upper troposphere, allowing urban air pollution to spread over much wider geographic areas. A typical reaction scheme for the production of PAN is presented in Table 4.
Fixed nitrogen is removed from the atmosphere by solution in rainwater or snow and by deposition on surfaces. The lifetime of NOx is relatively short—a month or so on average. Investigators have observed considerable temporal and spatial variability in NOx, with concentrations as high as 100 parts per billion (ppb) in polluted urban environments and as low as 50 parts per trillion (ppt) at remote locations, as, for instance, in the central Pacific. Fixed nitrogen is removed from the atmosphere mainly as HNO3, contributing to the acidity of precipitation. This matter will be discussed at greater length below in connection with the phenomenon of acid rain.

Carbon monoxide



Figure 9: Oxidation path for methane, CH4.
Photochemical processes play a major role not only for NOx but also for CO. Carbon monoxide is released to the atmosphere as a product of incomplete combustion. It also is formed as an intermediate in the oxidation of hydrocarbons, including methane. Present understanding of the path for decomposition of CH4 is shown in Figure 9. The hydroxyl radical, OH, is central to this process. Reaction with OH is the first step in the removal of many hydrocarbons, including CH4. Production of CO by oxidation of CH4 proceeds at a rate proportional to the concentration of OH, as does the removal of CO, which is effected by the reaction
The radical OH is formed by the reaction of the metastable oxygen atom O(1D) with H2O,
hν + O3 → O(1D) + O2.(20)
Reactions with CH4 and CO are the primary paths for removal of OH.
This leads to an interesting situation. An increase in carbon monoxide production due to fossil fuel combustion or biomass burning is expected to cause a decrease in OH concentration, with a resulting increase in the lifetime of gases normally removed by reaction with OH. As a consequence, one may expect an increase in CO concentration that is even larger than that due to growth in the magnitude of the source alone, with a related rise in the concentration of CH4. The increases in the global abundances of CO and CH4 that have been observed may be attributed, at least in part, to the complexities and nonlinearities of the chemistry affecting the concentration of OH.
Carbon monoxide has a lifetime in the atmosphere of about a month. The concentration in the Northern Hemisphere is larger than in the Southern Hemisphere by about a factor of 2. Surface concentrations at unpolluted sites range from 100 to 200 ppb and vary seasonally, with the highest values in winter when the OH concentration is low. The relative abundance, or mixing ratio, is large near the surface, decreasing aloft in the Northern Hemisphere, with an opposite trend in the south. Mixing ratios of CO in the Southern Hemisphere average about 60 ppb at the surface, climbing to roughly 70 ppb in the upper troposphere. Concentrations in cities reach values as high as one part per million and can pose a problem for public health from time to time.

Suite of sulfur compounds
In addition to the carbon, nitrogen, and oxygen compounds discussed above, the atmosphere contains a suite of sulfur-bearing gases. Like nitrogen, sulfur can occur in a variety of oxidation states. Compounds observed in the atmosphere range from the reduced gases carbonyl sulfide (COS), carbon disulfide (CS2), hydrogen sulfide (H2S), and dimethyl sulfide [(CH)2S], where sulfur is present in the oxidation state +2, to the more oxidized form of sulfur dioxide (SO2), where sulfur is in the oxidation state +4. The reduced species H2S and (CH3)2S are produced by reactions involving living organisms. Hydrogen sulfide originates mainly from anaerobic environments on land (e.g., swamps and marshes). The ocean is an important source of (CH3)2S. The origin of COS and CS2 is less clear; it may include contributions from a variety of sources, both natural and anthropogenic.
Reduced sulfur gases are oxidized at varying rates in the atmosphere and contribute to the global budget of SO2. The first step in oxidation for most involves reaction with the OH radical. Lifetimes of H2S and (CH3)2S are relatively brief, and the distribution of these species in the atmosphere is correspondingly variable. Carbonyl sulfide is longer lived, and this species is distributed more or less uniformly with latitude, with a mixing ratio of about 500 ppt. The lifetime of carbon disulfide is intermediate; its mixing ratio varies from about 3 to 30 ppt.
Sulfur dioxide is formed not only by oxidation of reduced gases in the atmosphere but also as a by-product of the combustion of coal and the smelting of ores such as copper. The anthropogenic component is comparable to the natural one on a global scale and is dominant over large areas of the industrially developed Northern Hemisphere. The lifetime of SO2 is about a week. Sulfur in SO2 is oxidized rapidly by both homogeneous and heterogeneous reactions and is removed from the atmosphere by precipitation and by dry deposition on surfaces, mainly as sulfuric acid (H2SO4) in which sulfur is present with oxidation state +6. The abundance of SO2 is variable, as would be expected given its relatively short lifetime. On a global basis, the mixing ratio averages about 200 ppt.

Suite of noble gases



Table 3 includes, in addition to the carbon, hydrogen, nitrogen, oxygen, and sulfur species discussed above, a suite of noble gases, with mixing ratios ranging from almost 1 percent in the case of argon to about 1 ppb for xenon. There are two components to the Earth's noble gas inventory. Some of the noble gases are primitive in the sense that the corresponding elements were captured from the solar nebula when the Earth formed some 4,500,000,000 years ago. Others, notably helium-4 and argon-40, are of more recent origin, produced by the radioactive decay of such elements as uranium, thorium, and potassium in the Earth's crust and mantle. Noble gases are released from the interior of the planet at ocean-spreading centres (i.e., the mid-oceanic ridges) when new crust is formed, as well as from volcanoes and hot springs. With the exception of helium, which escapes, the noble gases are inert and accumulate in the atmosphere. Studies of the elemental and isotopic composition of the atmospheric gases, in combination with analogous data for meteorites and other planets, provide valuable clues to the origin and early history of the Earth.

Concentrations of halogenated hydrocarbons



The atmosphere contains a variety of halogenated hydrocarbons as summarized in Table 3. These compounds are for the most part industrial in origin, with the exception of methyl chloride (CH3Cl) and methyl bromide (CH3Br), which are produced at least in part by marine organisms. Halogenated hydrocarbons are used as solvents, foam-blowing agents, fire extinguishers, fumigants, refrigerants, and propellants in aerosol spray dispensers. They are among the most ubiquitous products of modern industrial society.
Concentrations of several of the more widely used chlorinated species, CCl2F2 (CFC-12, or F-12), CCl3F (CFC-11, or F-11), CHClF2 (CFC-22), CH3CCl3 (methyl chloroform), and C2Cl3F3 (CFC-113), have been increasing at compound rates of between 5 and 10 percent over the past several decades. The fully halogenated methanes, CCl2F2 and CCl3F, are exceptionally stable, with lifetimes in the range of 60 to 130 years. They are removed mainly in the stratosphere by photolysis in the ultraviolet. They break down, releasing their constituent chlorine atoms to form reactive radicals, such as Cl and ClO. Chlorine radicals play an important role in the chemistry of stratospheric ozone (see below). There is concern that an increasing burden of stratospheric chlorine could result in a significant global reduction in the column abundance of ozone with an attendant increase in the transmission of ultraviolet solar radiation.
Fluorine is relatively inert in the stratosphere. Atoms of fluorine formed by the decomposition of the halocarbons react with methane to form hydrogen fluoride (HF). In contrast to hydrogen chloride (HCl), which decomposes rapidly by reaction with the OH radical, HF is stable in the stratosphere. The concentration of reactive fluorine radicals is much less than that of chlorine. Consequently, fluorine plays a negligible role in the chemistry of stratospheric ozone.
The abundance of brominated halocarbons is much less than the abundance of the chlorinated species. Hydrogen bromine, however, is less stable than HCl. A relatively large fraction of the bromine in brominated compounds that decompose in the stratosphere ends up in the form of the radicals Br and BrO. It is thought that these species can contribute significantly to removal of ozone. They tend to amplify the effect of chlorine and may become even more significant in the future.

The role of photochemistry
The preceding section dealt chiefly with gases that enter the atmosphere from below. As explained, the three primary sources for such gases are reactions effected by living organisms excluding man, reactions specifically attributable to human activity, and processes relating to the abiotic metabolism of the Earth. In this section, attention is focused on the role of photochemistry, the synthesis of gases in the atmosphere as a consequence of the interaction with sunlight. Ozone is by far the most important product of the photochemical process.
The bulk of the atmosphere's O3 is found in the stratosphere at altitudes of about 25 kilometres. It is this reservoir that is primarily responsible for protecting the surface from otherwise harmful rays of ultraviolet sunlight. The abundance of O3 in the troposphere also is environmentally significant. Direct exposure to high concentrations of O3 can reduce productivity in plants. It also can have a variety of undesirable effects on public health, causing, for example, respiratory problems.
The processes regulating the abundance of O3 in the stratosphere and troposphere are quite distinct. In the former case, production of O3 is initiated by photolysis of O2; in the latter, the O2 bond is broken through a complex set of reactions associated with the oxidation of hydrocarbons in the presence of NOx.

Stratospheric ozone
Studies of O3 have a long history, dating to about 1930. Early research emphasized the chemistry of a hypothetical pure oxygen system. Four reactions were considered:
and

Special Comp


Special Comp


Special Comp


Special Comp


Special Comp


Special Comp


Special Comp


Special Comp





Figure 10: Schematic view of ozone chemistry in a pure oxygen environment.
Reactions (22) and (23) serve to establish the relative abundances of atomic oxygen and ozone. The oxygen atom formed in reaction (21) cycles many times through (22) and (23) before it is eventually removed by (24). It is thus convenient to consider O and O3 in combination, to think of a family of odd oxygen compounds, O + O3. Reaction (21) is the source of odd oxygen, and reaction (24) is the sink. In equilibrium, source and sink must balance. This simple scheme is illustrated in Figure 10.
The concentration of O is controlled by reactions (22) and (23) at all levels below about 70 kilometres. Thus,
Balancing (21) and (24), using (25), gives
where J and k denote rate constants for (21,22,23,)–(24) with the subscripts identifying the relevant reactions; fO2 is the mixing ratio for O2, [O2] = fO2 [M]. (Concentration is signified by the bracketing of symbols.)

Special Comp





Figure 11: Height profiles for O3, O, and M illustrating important terms in the equation …
The essential features of the photochemistry of stratospheric O3 are summarized by reaction (26). The rate for photolysis of O2, J21, is vanishingly small at low altitude: photons with sufficient energy to dissociate O2 are absorbed by O2 and O3 above 25 kilometres. Rate constants are more or less invariant with height at high altitude, and the abundance of O3 decreases accordingly with altitude, as [M]3/2 at high elevation. The concentration of O3 is small at low altitude, limited by the small and declining value of J21. It follows that O3 must assume a maximum value at some intermediate level. In practice, the maximum is observed at a height of about 25 kilometres. Height variations of J21, J23, [M], [O], and [O3] are shown schematically in Figure 11.
Much of the complexity of stratospheric chemistry, and indeed much of its recent history, is concerned with working out the paths for removal of odd oxygen besides reaction (24). For example, the pair of reactions
have a net effect equivalent to (24). In the presence of nitric oxide, NO, the rate for (24) is faster than would be otherwise the case; in short, reaction (24) is catalyzed by NO. Similarly, (24) can be catalyzed by chlorine, Cl,
by bromine, Br,
or by the hydroxyl radical, OH,
Catalytic paths occur for which there is no direct reaction analogue. The sequence
is equivalent to
as is
The rate constant for (37) is vanishingly small, and this reaction has a negligible effect on stratospheric ozone. The catalytic paths, (35) + (36) in particular, can be important, however, and provide additional means for the removal of odd oxygen, especially at low altitude where the abundance of atomic oxygen is small.
It is convenient to extend the definition of odd oxygen to include NO2, ClO, BrO, and HO2 (hydroperoxyl). The catalog of sinks for odd oxygen expands then to encompass reactions (28), (30), (32), (34), and (40). Note that each of these loss reactions is responsible for the removal of two compounds of odd oxygen. Similarly, (21), the only source reaction, is responsible for two odd oxygen products.

Significance of the concept of odd oxygen
The concept of odd oxygen is useful in that it provides a systematic scheme for classifying the several hundred reactions involved in a comprehensive model of stratospheric O3. Only a few of these reactions are directly implicated in the removal of O3. The balance is important mainly in regulating the levels of NO2, ClO, BrO, and HO2.
It is useful to associate these compounds with specialized families in addition to odd oxygen. The choice of association is to some extent arbitrary. The family of NOx compounds mentioned earlier provides a natural link for NO2. In similar fashion, ClO is identified with Clx, Cl + ClO + ClNO3 + HOCl + HCl; BrO with Brx, Br + BrO + BrNO3 + HBr; and HO2 with HOx, OH + HO2 + H2O2. It is important to note that a particular compound may belong to more than one family.
Reaction of the metastable oxygen atom O(1D) with nitrous oxide, N2O,
is the primary source of stratospheric NOx. Decomposition of the halocarbons provides the dominant source of Clx and Brx, while reaction
is the most important source of HOx. Transport to the troposphere is the major means for removal of NOx, Clx, and Brx from the stratosphere. HOx is removed in situ by
and by




Figure 12: Transformations of nitrogen compounds in the atmosphere.



Figure 13: Transformations of chlorine compounds in the atmosphere.



Figure 14: Transformations of bromine compounds in the atmosphere.



Figure 15: Transformations of hydrogen compounds in the atmosphere.
Sources, sinks, and interchange mechanisms are illustrated schematically for NOx, Clx, Brx, and HOx in Figures 12, 13, 14, and15.

Distribution of ozone



Figure 16: Column abundance of O3 as function of latitude and time of year in Dobson …
The distribution of O3 with altitude and latitude is determined not only by chemistry but also by transport. The lifetime of odd oxygen, mainly O3, is short at high altitude and low latitude where the abundance of atomic oxygen is relatively high. Chemistry is the dominant control in this case. The influence of transport is important at low altitude and high latitude; O3 tends to accumulate in these regions where it is protected from chemical loss. This accounts for the gross features of the latitudinal distribution displayed in Figure 16. There is a net flow of O3 from the stratosphere to the troposphere, where it is removed heterogeneously at the surface by gas phase reaction with HO2.

Tropospheric ozone
Tropospheric O3 is supplied either by in situ oxidation of hydrocarbons and CO or by downward transport from the stratosphere. The path for production due to oxidation of CO is particularly simple and may be summarized as follows:
and
An alternate path for oxidation of CO involves reactions (45) and (46) followed by
The bulk reaction in this case is
In the first path O3 is produced, while in the second it is consumed. The choice of oxidation scheme depends on the relative abundance of NO (nitric oxide) and O3, on the relative importance of reactions (47) and (51). Production of O3 is favoured if the abundance of NO is high; loss is predominant if NO is low. Similar bifurcation occurs in the oxidation of CH4 and other more complex hydrocarbons.
The chemistry of tropospheric O3 is in general less well understood than the chemistry of stratospheric O3, because the troposphere is intrinsically more complex than the stratosphere. Distributions of NOx and hydrocarbons are quite inhomogeneous and, as yet, not well defined. Ozone can be produced in one region and consumed in another. In any case, further research is required to understand better the intricacies of tropospheric chemistry.

Effects of human activity on atmospheric composition and their ramifications
Change has been a fact of life for the Earth for most of its history. Changes in the past have been both fast and slow, regular and episodic, driven by factors internal to the planet and by forces from without. The evolution of an oxygen-mediating enzyme altered the composition of the atmosphere irreversibly some 2,000,000,000 years ago, allowing free oxygen to accumulate and thereby changing the nature of the biota for all subsequent time. The impact of a large asteroid (or possibly a comet), it is thought, led to the demise of the dinosaurs and a variety of other species at the end of the Cretaceous Period (from about 136,000,000 to 65,000,000 years ago). Climate has adjusted throughout geologic time to the changing position of the continents driven by the slow motion of giant lithospheric plates, rigid blocks of the Earth's surface rocks (see plate tectonics). Ice sheets have waxed and waned with regularity in response to the changing orbit and rotation axis of the Earth, and, as noted above, there have been associated changes in the carbon dioxide level of the atmosphere. What sets the present era apart, however, is the unique importance of a single species, man. Since about the mid-19th century, humankind has developed the capacity to harness the powers of nature, to mine the reserves of organic carbon laid down in sediments over millions of years. It has used this abundant source of energy to alter its way of life, with effects, for the most part inadvertent though undoubtedly large, for the global environment. Some of those effects associated with the atmosphere are considered here.

Climate modification
Trace gases, carbon dioxide and water vapour in particular, play an essential role in terrestrial climate. In the absence of these molecules that absorb strongly in the infrared, the surface temperature would be about 40 K colder than it is today. The oceans would be frozen over, and life as we know it would be impossible. There is an important synergism between CO2 and H2O. Carbon dioxide itself absorbs only a small fraction of the heat radiated by the Earth's surface. Water vapour is much more significant. The abundance of H2O, however, is controlled ultimately by temperature. An increase in CO2 may be expected to cause an increase in temperature, which allows more water vapour to enter the atmosphere and leads to a further increase in temperature.
It has been suggested that the climate of the Earth could be ultimately unstable. Addition of gases capable of trapping heat could accelerate the release of H2O and raise the temperature to a point where the oceans would evaporate, transforming the atmosphere into a hot steamy vapour. The surface pressure of the atmosphere would then be about 1,000 times higher than it is today, and its dominant constituent would be H2O. Some believe that such changes may have occurred on Venus, with H2O subsequently escaping from the top of its atmosphere. In any event, contemporary Venus is a striking example of the importance of the greenhouse effect. Its atmosphere contains a large concentration of CO2, almost 106 times more than that of the Earth. Moreover, the Venusian surface temperature is much hotter than the Earth's—about 780 K—in spite of the fact that Venus absorbs less energy from the Sun because of its ubiquitous cloud cover and associated high albedo.

Increase in carbon dioxide concentration
The concentration of CO2 in the Earth's atmosphere has risen steadily over the past 140 or so years, from about 280 parts per million in 1850 to about 350 parts per million. The change is due largely to the combustion of fossil fuels. Since the Industrial Revolution, nearly 1.5 × 1011 metric tons of organic carbon have been mined and consumed in the form of coal, oil, and natural gas. Carbon dioxide is the largest single waste product of modern society. The average person is responsible for the release of almost four tons of CO2 each year; the amount is even larger in the developed countries. The total global emission in 1985 was close to 5 × 109 metric tons of C.
Approximately half of the carbon emitted since the Industrial Revolution persists in the atmosphere today. The balance is presumed to have made its way into the oceans or to have been incorporated into organic matter on land. Uptake of CO2 by the oceans is limited by the supply of carbonate ions, CO32-, in surface waters. The bulk reaction is
Carbon dioxide is a weak acid, and inclusion of CO2 in the oceans leads to a reduction in pH, transforming carbon from CO32- to bicarbonate ions, HCO3-, and dissolved neutral carbon.There is a limit, however, to this switching. The total negative charge carried by dissolved carbon compounds is fixed by the alkalinity of the ocean waters:
Alkalinity can be altered only by the addition of salts, which are supplied, for example, by the dissolution of calcite, CaCO3, in sediments. The CO32- content of waters at the ocean surface is small, and sustained uptake of CO2 requires that CO32- be supplied continuously to the surface. Over a 100-year period about 10 percent of the water in the oceans is exposed at the surface. If such is the case, only 30 percent of the carbon dioxide released to the atmosphere by burning fossil fuels can be incorporated into the sea.
A continuing rise in carbon dioxide is inevitable. If current estimates for the reserve of fossil fuels—about 4 × 1012 metric tons of carbon—are considered, and if it is assumed that half of this reserve is used up over the next 100 years, the level of CO2 could exceed 1,000 parts per million. If one assumes more conservatively that the consumption of fossil fuels will double over the next 100 years, CO2 may be expected to grow to about 600 parts per million—approximately twice the value in 1850.

Impact of other gases on climate
Carbon dioxide is not the only gas with the ability to affect climate. The trace gases F-11, F-12, methane, and nitrous oxide are even more potent on a molecule-per-molecule basis. It has been estimated that a molecule of F-11 or F-12, for example, can produce warming equivalent to that caused by 104 molecules of CO2. Perturbations to stratospheric ozone and tropospheric ozone also can have a significant impact. In short, a comprehensive theory of climate change must allow for all of these various changes in atmospheric composition.
Most assessments of climatic change are based on simple, one-dimensional energy-balance models. These models, it is hoped, will provide an estimate for the globally averaged change in surface temperature. Studies suggest that the rise in carbon dioxide since 1850 should have resulted in an increase in the contemporary surface temperature of about 0.4° C. The trace gases F-11, F-12, methane, nitrous oxide, ozone, and stratospheric water vapour are estimated to have added to heating by CO2 an increment equivalent of an additional 0.16° C, for a net change of 0.56° C. A change of this magnitude is generally consistent with observation.




Figure 17: Changes in the global mean temperature expected due to past and anticipated variations …
The effects of trace gases are expected to grow relative to CO2 in the years ahead. Results from one particular simulation are given in Figure 17. In this case, CO2 was predicted to cause an increase in surface temperature of about 0.9° C by the year 2030. With other trace gases included, it was estimated that the temperature increase could be as large as 2.2° C. A change of this magnitude would be unprecedented in recent Earth history.
A prediction of an increase in global mean temperature, whatever its magnitude, has at best limited use for government policymakers. Scientific concern is more with the geographic and temporal distribution of the change. Will it rain more or less, and when? How will the change in climate affect food production? What changes are to be expected in the location of deserts, forests, grasslands, and agricultural zones? Benefits may be expected for some and problems for others. Comprehensive assessment requires a model incorporating all of the important feedbacks between the atmosphere, oceans, and biosphere, with spatial resolution sufficient to distinguish at least the major biomes. Such a model is as yet unavailable, though efforts to develop it are under way. Preliminary studies with three-dimensional models similar to those used for predicting weather suggest that climatic zones can be expected to shift to higher latitudes on a warmer Earth. Further work is required, however, before models of this kind can be used with confidence for quantitative planning.

Depletion of stratospheric ozone
The importance of stratospheric O3 has been recognized in a general way for almost 50 years. In the absence of O3, the surface of the Earth would be exposed to lethal ultraviolet radiation with wavelengths as short as 240 nanometres. It was only in 1970, however, that scientists began to focus on the fact that even small changes in O3 can have a significant impact on humans. Investigators observed that migration of people to lower latitudes—the shift in population from the northeastern part of the United States to the Sun Belt (roughly the southern and southwestern regions of the country), for example—was accompanied by an alarming rise in the incidence of skin cancer. Not all of this increase could be attributed to enhanced sunlight. There appeared to be an underlying factor to the smaller abundance of O3 at lower latitudes and the associated increase in exposure of fair-skinned people to solar radiation with wavelengths near 300 nanometres. Epidemiological studies suggested that the incidence of skin cancer would rise by about 3x percent for every x percent decrease in the column density of O3. This led to inevitable questions concerning the stability of O3 as the human influence began to extend upward through the tropopause to the stratosphere.
It has long been known that the stratosphere turns over very slowly. Debris from the testing of nuclear bombs in the late 1950s and early 1960s was readily detectable a decade later. Plans were under way in the early 1970s to develop a commercial fleet of supersonic aircraft. These planes were projected to cruise at altitudes of about 20 kilometres. It seemed inevitable that gases from their exhaust would accumulate in the stratosphere, and nitric oxide became a particular concern. It was suggested that nitric oxide from a fleet of 500 supersonic aircraft could lead to a reduction in ozone abundance by as much as 3 percent. This led to a major research program coordinated by the U.S. Department of Transportation, and results from the program played an instrumental role in the decision by the federal government to suspend funds for the development of the large supersonic transport (SST). The British and French, however, continued their work on a smaller version of a supersonic transport, the Concorde, which was eventually introduced for limited service from Europe to North America and the Middle East. Flying at a lower altitude than the SST would have and releasing smaller quantities of nitric oxide, the Concorde has had a negligible effect on stratospheric ozone. It remains as a reminder of a vitriolic debate that served, if nothing else, to draw attention to the stratosphere, to highlight the potential vulnerability of even the most remote regions of the atmospheric environment.
The debate concerning the environmental impact of the SST has had a lasting effect on the development of atmospheric science, spawning a new interdisciplinary program of research linking chemists, physicists, and biologists in a common effort to understand the stratosphere. The program, with international participation, has been remarkably successful and has led to a new view of the interdependence of the atmosphere, hydrosphere, and biosphere.
The concern over the effects of exhaust gases from supersonic aircraft was soon followed by a new issue: the possibility that chlorine atoms released by decomposition of chlorofluorocarbons could have a larger and more persistent effect on stratospheric ozone. CFC's were developed first in the 1930s but found widespread use only in the years following World War II. They were employed with great success by U.S. troops in the Pacific to dispense insecticides from aerosol spray cans. This led to many commercial uses, from propellants and refrigerants to foaming agents and degreasers and a host of other applications. Moreover, the use of CFC's rapidly spread from the United States to Europe and the Far East. All this changed in 1975, when it was recognized that the release of CFC's to the atmosphere could pose a serious problem for stratospheric ozone. Production of F-12 declined from a peak of about 4.5 × 105 metric tons in 1975 to about 3.4 × 105 metric tons in 1982. A similar drop was registered for F-11.
Much of the work undertaken since the mid-1970s has focused on the effects of CFC's on the assumption that the composition of the atmosphere was otherwise constant. It has become clear, however, that the response of ozone depends not simply on the abundance of CFC's but also on the abundances of methane, nitrous oxide, and carbon monoxide. These species, too, are changing. Current models suggest that a continuing release of CFC's at the rate registered in 1980, other gases remaining constant, would lead to a reduction in stratospheric ozone by about 5 percent. Maximum impact is predicted to occur at altitudes above 25 kilometres. An increase in nitrous oxide of 20 percent is expected to cause a reduction in ozone of about 2 percent. An increase in carbon dioxide should lead to a reduction in stratospheric temperatures with a consequent reduction in the anticipated impact of CFC's and nitrous oxide on ozone. The effect of an increasing burden of methane is more complex. Oxidation of methane provides a source of ozone at low altitude, while the reaction of chlorine with methane converts chlorine radicals to hydrogen chloride, resulting in a reduction in the impact of CFC's between 30 and 40 kilometres.
Models suggest that the change in the column density of stratospheric ozone to date should be relatively small. Reductions in ozone at high altitude, near 40 kilometres, ought to be balanced by excess production at low altitudes due in part to the higher level of methane and in part to NOx released by high-altitude aircraft. Observational evidence is consistent with this view. A statistical analysis concluded that the change in the ozone column from 1970 to 1983 averaged -0.003 percent per decade. It showed, however, that a small though statistically significant drop in ozone—a decline of about 2 percent—occurred at altitudes above 30 kilometres between 1970 and 1980.
There has been a new development since 1985. That year, Joseph C. Farman and his associates at the British Antarctic Survey reported that the level of ozone over Antarctica had dropped precipitously every October since 1982, with the first such change apparent as early as 1978. A number of theories, or more properly hypotheses, have been advanced to account for this phenomenon. Several implicate effects of anthropogenic chlorine, enhanced by small quantities of bromine. Others suggest that the reduction may be due to a diminished supply of ozone from low latitudes, reflective of a change in stratospheric dynamics. In any case, the phenomenon was quite unexpected and serves as a powerful warning that current scientific understanding of the stratosphere is still rudimentary.

Acid rain and allied problems
The acidity of rain, or of any liquid, depends on the concentration of protons (positively charged hydrogen ions), H+, in solution. It is measured conventionally by pH, defined by the relation
where [H+] denotes the molar abundance of H+—i.e., the concentration expressed in moles per litre. Pure water is characterized by a pH of 7.0. A liquid is said to be acidic if pH is less than 7.0; otherwise it is considered neutral or alkaline.
Exposed to an atmosphere containing carbon dioxide, water absorbs carbon, releasing protons by
Similar reduction in pH occurs with absorption of nitric acid (HNO3) or sulfuric acid (H2SO4). Reactions in this case may be represented by
and
Absorption of less oxidized gases such as nitrogen dioxide (NO2) or sulfur dioxide (SO2) may be described by
and
In each case, protons are liberated and pH declines. This result may be viewed as an inevitable consequence of the chemical nature of carbon, nitrogen, and sulfur. These elements are poised to form negative ions in their oxidized states. The resulting negative charge must be balanced, to some extent at least, by production of H+.
The pH of rain in equilibrium with atmospheric CO2 would have a value of 5.65. The addition of HNO3 SO2 with concentrations typical of the clean environment remote from concentrated sources, about 100 parts per trillion, would lower the pH to approximately 5.3. (This assumes lifetimes for nitric acid and sulfur dioxide of roughly 10 days with precipitation at an average rate of one metre per year.) A concentration of sea salt or alkaline aerosol equivalent to about 50 parts per million, a typical value, would raise the pH to about 6.0. On this basis, it is concluded that the natural range of pH should lie between 5.3 and about 6.0. The data in Table 5 are consistent with this view, except for the measurement in Brazil indicating a value of 4.6. The acid content of rain in the Amazon River basin may reflect the importance of local biologic sources of sulfur, particularly hydrogen sulfide (H2S). As mentioned above, hydrogen sulfide is oxidized readily in the atmosphere to sulfur dioxide, with a lifetime of no more than a few days.
Rains in the northeastern part of the United States show median values for pH in the range of 4.1 to about 4.5. Values are lowest in a band extending from the Ohio River Valley to upstate regions of New York. Information on the composition of precipitation is particularly illuminating. Table 6 shows data from sites in Minnesota, Georgia, and New York obtained in 1979. Sulfate accounts for about 60 percent of the negative charge in rains at all three sites; nitrate ranges from 20 percent in Georgia to 34 percent in New York. The pH in New York was 4.34, as compared with the high value, 6.31, measured in Minnesota. The difference may be attributed to the presence of alkaline aerosol in rain reaching the site at Minnesota. The addition of calcium carbonate, CaCO3, results in removal of H+ according to
Soils in the western part of the United States are generally alkaline, and the aerosol content of the atmosphere tends to reflect this condition. Soils in the eastern United States are more acidic, and the buffering capacity is reduced accordingly.

Sources of acid-forming compounds
The low pH and high sulfate content of precipitation in the Northeast is scarcely surprising. Approximately 12,000,000 metric tons of sulfur were vented to the atmosphere from industrial sources east of the Mississippi River in 1978. The sulfur content of rain falling on the same region amounted to only 20 percent of the quantity released. The balance was either exported or deposited by dry processes on various surfaces—vegetation and so forth. The industrial source grew by a factor of 2.5 between 1950 and 1970, with the rate of growth most rapid during the 1950s. The decline in subsequent growth reflected a shift by utilities from coal to oil and nuclear power and from high-sulfur to low-sulfur coal, a trend that has been reversed somewhat since the early 1980s. The global industrial output of sulfur in 1978 was about 65,000,000 metric tons, which may be compared with estimates of 64,000,000 metric tons per year for production by the unperturbed global biosphere. The biospheric production in the eastern United States is about 100,000 metric tons per year and is surely less than the industrial contribution by at least a factor of 10.
Industrial sources of oxidized nitrogen are similarly dominant. The global industrial output of NOx (see above) is about 20,000,000 metric tons N per year, distributed more or less equally between North America, Europe, and the rest of the world. The industrial output is thought to exceed production by natural agents in North America by a factor of from 3 to 13. Combustion of fossil fuels accounts for about 40 percent of the NOx. Biomass burning accounts for another 25 percent, with the remainder from lightning and microbial processes in soils.
The relation between emission of potentially acidic compounds and ultimate delivery of acids to specific geographic regions is by no means straightforward. It depends on the nature of the atmospheric circulation and characteristics of the source—e.g., height of smokestacks. Chemical factors controlling dry deposition are poorly understood, and there are gaps in scientific understanding of the relevant aspects of atmospheric chemistry. Oxidation of nitrogen dioxide to nitric acid is thought to proceed by the reaction
Oxidation of sulfur is initiated similarly by reaction with the hydroxyl radical, OH,
or it may proceed heterogeneously with hydrogen peroxide, H2O2, formed from HO2. Ammonia, NH3, is important in neutralizing acid aerosols. A complete description of the relevant chemistry must thus include NH3 and O3 in addition to oxides of nitrogen and sulfur. Interconnections between various chemical cycles can lead to subtle, unexpected nonlinearities in the composite effects, with important consequences for the geographic pattern of the resulting deposition of acid.

Effects of acid rain
Acids can affect receiving systems in a variety of ways, some good, others bad. Addition of nitrate to ecosystems deficient in nitrogen can lead to enhanced biologic productivity. The benefits, however, may be offset by toxic effects associated with the release of aluminum and magnesium from clay minerals if the pH of soils falls below about 5.5. Leaching of metals from soils and sediments also can have adverse effects on aquatic systems. Fish may be killed by aluminum even under conditions where the pH is considered otherwise benign. Given the current rates of acid deposition, the pH of a number of lakes in the northeastern United States and in southeastern Canada is expected to average 4.2 to 4.8, with occasional depressions to near 3.8 during periods of snowmelt or heavy rain. Most fish species and almost all mollusks perish if the pH falls below 4.8; planktonic communities are simplified in this case, dominated by a few acid-tolerant taxa. Adverse effects of acid precipitation are most obvious for lakes, and it is not surprising that this aspect of the issue has received the most attention in the press. The effects on the terrestrial biosphere are more ambiguous. It appears that elevated levels of ozone in the troposphere may pose a more serious hazard for plant life even at locations removed from major industrial sites. Effects in this case, for a variety of agronomic and forest species, include foliar injury and significant reduction in growth and yield.
In summary, there are deficiencies in current scientific understanding of acid rain, both in its atmospheric and biospheric dimensions. Uncertainties can be removed or reduced by a suitably directed research program. It is clear that the chemistry of the troposphere, both polluted and clean, should receive attention in this effort.

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 nm—mainly from photons in the strong solar emission line at Lyman α, 121.6 nm. Lyman α 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 α is able to provide a significant source of ionization for the D region near 90 kilometres.

Ionospheric variations
The ionosphere is variable in space and time. Some of the changes are chemical in origin and can be readily understood on the basis of the general considerations outlined above. There is a systematic variation, for example, as a function of time of day. In the early morning the Sun is relatively low in the sky. Radiation must penetrate a large column of air before reaching a given level of the atmosphere. Ionization rates are less than at noon, and the profile of ionization is shifted to higher altitudes. The heights of the D, E, and F1 layers change in response to solar elevation. The layers are lowest and the densities of electrons are highest at noon. Ionization in the D, E, and F1 regions tends to disappear at night as a result of recombination in the absence of a compensating source.
The diurnal variation of the F2 layer is less dramatic. Ionization produced at high altitudes during the day maintains a sizable density of electrons at the F2 peak throughout the day and diffuses downward at night. This accounts for the fact that radio reception (both in the broadcast and shortwave bands) is generally best at night. Ionization at lower altitudes—primarily those corresponding to the D region—tends to interfere with radio transmissions during the day. Interference is minimal at night, since ionization in the D layer effectively disappears with the setting of the Sun.
The density of ionization varies as a function of solar activity in response to changes in the intensity and spectral properties of solar radiation. The output of solar energy is relatively constant in the visible and near-ultraviolet portions of the spectrum. It varies appreciably, however, at shorter wavelengths, reflecting changes in the temperature of the outermost regions of the solar atmosphere. The changes are particularly large, in excess of a factor of 10, at X-ray wavelengths. Variations in the D region are correspondingly large, with smaller though still significant changes in the E and F layers. The temperature of the thermosphere varies as a function of solar activity in response to the changing input of energy. Exospheric temperatures range from about 750 K at solar minimum to as high as 2,000 K when activity is particularly intense.
Solar activity varies on a characteristic time scale of about 11 years. It is not entirely periodic, however, as successive cycles can differ significantly. There are indications that activity can be low for periods as long as centuries. For example, the Sun was quiet for more than 200 years from about 1600 to roughly 1850.
As mentioned earlier, ionization above the F2 peak is removed mainly by downward diffusion. Ionization is constrained, however, to move along the magnetic field. The field is oriented horizontally at the magnetic equator, so vertical diffusion is inhibited at low latitudes. The density of O+ and electrons at low latitudes is controlled by chemistry to a larger extent than at high latitudes. The F2 peak is correspondingly higher in altitude and the density of electrons is elevated accordingly.
Ions and electrons formed at high altitudes and low latitudes are transported to higher latitudes by thermospheric winds. As a result, the highest density of electrons at the F2 peak is observed at intermediate latitudes, displaced by about 10° relative to the magnetic equator.
Transport also can affect the distribution of ionization at lower altitudes. The diurnal pattern of heating in the troposphere and stratosphere excites a spectrum of waves, some of which are free to propagate vertically. The amplitude of the waves grows significantly as the disturbance enters regions of lower density. The passage of the waves is associated with strong alternating horizontal winds. Ionization can be driven up inclined magnetic field lines at one altitude, while winds blowing in an opposite direction at higher altitudes can induce simultaneous downward motion. This can lead to bunching of ionization, causing a local enhancement of the electron density. The mechanism is particularly important in the E region and is responsible for a phenomenon known as sporadic E.
The buildup of ionization is normally limited by dissociative recombination of molecular ions. The ionosphere at D- and E-region altitudes, however, contains a small but variable concentration of atomic ions derived from the ionization of metals ablated from meteorites. The density of metallic ions, notably those of sodium, magnesium, and potassium is sometimes high enough to supply, after concentration, a layer of ionization with density comparable to that of the F layer. This can result in a major temporary disruption of radio communications.
Winds generated in the lower ionosphere by thermal forcing from below have characteristic periods expressed as submultiples of a day. Waves with a period of 24 hours dominate at low latitudes, whereas those with a characteristic period of 12 hours are more important at high latitudes. The waves are basically similar in physical origin to tides excited in the ocean by lunar gravity. The vertical motion to which they owe their origin is generated, however, by the diurnal pattern of heating and cooling rather than by gravity. Additional waves can arise due to irregular forcing associated, for example, with thunderstorms, motion over orographic features, and a range of other small-scale meteorologic disturbances. These small-scale disturbances are referred to as gravity waves to distinguish them from the more regular planetary-scale motions excited by the diurnal cycle of heating and cooling. The regular response to thermal forcing is known as the atmospheric tide.
Tides and gravity waves have similar effects on ionization in the E region. They both are responsible for concentrating ionization in layers. In combination with the large-scale system of winds in the lower thermosphere, they also are effective in driving an irregular current that flows in the E and lower F regions of the ionosphere. The current owes its origin to differences in the facility with which motions of ions and electrons are constrained by the magnetic field. It is associated with an electric field and results in a modulation of the magnetic field that can be readily detected at the surface. The current is particularly intense in the equatorial region where it is known as the electrojet. The region of strong current flow is called the dynamo region.
Protons, H+, and helium ions, He+, are important components of the ionosphere above the F2 peak. They increase in abundance relative to the atomic ion of oxygen, O+, with increasing altitude. Protons are produced by the photoionization of atomic hydrogen,
and by charge transfer from O+ to H,
Helium ions are formed by the photoionization of helium. The distribution of H+ and He+ with altitude reflects the influence of the polarization electric field set up to preserve charge neutrality. When O+ is the dominant ion, the polarization field acts to lift H+ and He+ with a force equivalent, but in opposite direction, to that exerted by the gravitational field on a particle of mass 8 atomic units. Protons behave as though they had an effective gravitational mass of -7 atomic units (1 - 8). The effective mass of He+ is -4 atomic units (4 - 8).
The abundance of H+ and He+ increases with altitude. Eventually H+ becomes the dominant component of the ionosphere, and the polarization field is then diminished. A field equivalent to the gravitational force acting on a body of mass 0.5 atomic units, directed upward for ions and downward for electrons, is sufficient to maintain equal densities of H+ and electrons. The effective masses of O+ and He+ shift to 15.5 atomic units (16 - 0.5) and 3.5 atomic units (4 - 0.5), respectively, and densities of O+, He+, and H+ decline with further increases in altitude.
The overall structure of the outer ionosphere is influenced strongly by the configuration of the magnetic field. Ionization is removed by diffusion, but diffusion is limited to the directions imposed by the morphology of the field.
Close to the Earth, the magnetic field has a structure similar to that of an ideal dipole. Field lines are oriented more or less vertically at high latitudes, sweep back over the Equator when they are essentially horizontal, and connect to the Earth in a symmetrical pattern at high latitudes in the opposite hemisphere. The field departs from this ideal dipolar configuration, however, at high altitudes. There, the terrestrial field is distorted to a significant extent by the solar wind with its imbedded solar magnetic field. Ultimately the terrestrial field is dominated by the interplanetary field. The solar wind compresses the field of the Earth on the dayside at a distance of about 65,000 kilometres from the planet. The terrestrial field is stretched out on the nightside in a giant tail that reaches past the orbit of the Moon, extending perhaps to distances in excess of 1,000 Earth radii. The boundary between regions dominated by the solar and terrestrial magnetic fields is known as the magnetopause, and the region interior to the magnetopause is called the magnetosphere.
The outermost regions of the magnetosphere are exceedingly complex, especially at high latitudes where the terrestrial field lines are open. Ionization from the solar wind can leak into the magnetosphere in a number of ways. It can enter by turbulent exchange at the dayside magnetopause or more directly at cusps in the magnetopause at high latitudes, where closed loops of the magnetic field on the dayside meet fields connecting to the magnetotail. In addition, it can enter at large distances on the nightside where the magnetic pressure is relatively low and the field lines are able to reconnect readily, providing easy access to the giant plasma sheet in the interior of the magnetotail. For further information about the magnetosphere and related subjects, see Earth: The magnetic field of the Earth.

Auroras
Auroras are perhaps the most spectacular manifestations of the complex interaction of the solar wind with the outer atmosphere. The energetic electrons and protons responsible for auroras are thought to emanate from the plasma sheet inside the Earth's magnetotail.
Auroras occur in both hemispheres. They are confined for the most part to high latitudes, to oval-shaped regions maintaining a more or less fixed orientation with respect to the Sun. The centre of an auroral oval is displaced a few degrees to the nightside with respect to the geomagnetic pole. The midnight portion of an oval is, on average, at a geomagnetic latitude of 67°; the midday portion is at about 76°. An observer between 67° and 74° magnetic latitudes generally encounters auroras twice a day—once in the evening and once in the morning.

Auroral zone
The portion of the Earth that traverses the midnight portion of an oval is known as the auroral zone. In the Northern Hemisphere, for example, it lies along the curve extending from the northern regions of Scandinavia through Iceland, the southern tip of Greenland, the southern region of Hudson Bay, Alaska, and on to the coast of Siberia. This is the prime region from which to view an aurora. The phenomenon is by no means static, however. The auroral zone shifts poleward at times of low solar activity. It has been known to move as far south as 40° (geographic latitude) in the United States during periods of high solar activity. At low latitudes, an aurora assumes a characteristic red colour. In ancient times this feature was often interpreted as evidence of impending disaster. Auroras assume a variety of forms, depending on the vantage point from which they are observed. The luminosity of an aurora is generally aligned with the magnetic field. Field lines are close to vertical in polar regions, and so an aurora occurring there appears to stand on end, to hang from the sky in great luminous drapes. It is a spectacular sight indeed, especially if viewed from a distance either from the north or south. At lower latitudes, the magnetic field lines are inclined with respect to the vertical. There, an aurora appears as streamers radiating from the zenith. Such is the majesty of auroras that no two displays are totally alike. Light can move rapidly across the sky on some occasions, and at other times, it can appear to stand in place, flickering on and off.

Causes of auroral displays
The most common type of aurora is associated with the bombardment of the atmosphere by electrons having energies of up to 10,000 electron volts (eV). The energy for these electrons originates ultimately from the Sun. It is propagated through space by the solar wind and is communicated to the magnetosphere, most probably to the plasma sheet, by complex electromagnetic interactions not as yet well understood. The energetic electrons enter the atmosphere along magnetic field lines. They produce a shower of secondary and tertiary electrons, approximately one for every 35 eV of energy in the primary stream. Primaries can propagate to altitudes as low as 100 kilometres. Most of the luminosity is produced, however, by low-energy secondary and tertiary electrons. Prominent emissions are associated with the red line of atomic oxygen at 630 nm, the green line of atomic oxygen at 558 nm, the first negative bands of ionized molecular nitrogen at 391 nm and 428 nm, and a host of emissions from O, O2, O2+ and N2. Many of these features also are present in the day and night airglow. They are most notable in auroras because of their intensity and the rapidity with which they switch on and off in response to changes in the flux and energy of incoming primaries. An aurora has a characteristic red colour if the energy of primaries is relatively low. In this case, emission is dominated by atomic oxygen and is confined for the most part to altitudes above 200 kilometres. If the energy of the primaries is high, an aurora has a greenish blue colour and extends downward to altitudes as low as 90 kilometres.
Auroral displays also are produced by the bombardment of the atmosphere by energetic protons. Protons with energies of up to 200,000 eV are responsible for auroral activity in a diffuse belt shifted equatorward from the main auroral zone. These protons can be detected from the ground by observation of Doppler-shifted radiation emitted by fast hydrogen atoms formed by charge transfer from atmospheric atoms and molecules. Protons also play a role at higher latitudes, especially at times following major flares on the Sun. It is thought that the protons responsible for auroras at the polar caps are solar in origin. Associated energies may reach as high as 106 eV, and particles may penetrate as deep as 80 kilometres. Polar cap auroras can provide a significant transient source of mesospheric and stratospheric nitric oxide. They can be responsible for small but detectable short-term fluctuations in the abundance of stratospheric ozone.

The Van Allen radiation belts



Figure 18: The Van Allen radiation belts contained within the Earth's magnetosphere.
The magnetosphere includes two doughnut-shaped belts, or zones, centred on the Equator that are occupied by appreciable numbers of energetic protons and electrons trapped in the magnetic field high above the sensible atmosphere (Figure 18). The inner belt extends from roughly 1,000 to 5,000 kilometres above the terrestrial surface and the outer belt from some 15,000 to 25,000 kilometres. The belts were named in honour of James A. Van Allen, the American physicist who discovered them in 1958. Van Allen's find was a triumph of serendipity: he detected the presence of the trapped particles with a Geiger counter designed to measure the flux of cosmic rays in space. It was the first great discovery of the space age and was achieved by combining data obtained with instruments carried by three of the earliest U.S. scientific satellites—Explorer 1, Explorer 4, and Pioneer 3.
The flux of protons crossing a square centimetre of surface in the inner Van Allen belt can be as large as 20,000 per second, higher than the flux of cosmic radiation in space by a factor of 104. Protons in the inner belt have energies in excess of 7 × 108 electron volts, enough to enable them to penetrate several centimetres of lead. Spacecraft that have to fly through the belts thus must be protected; otherwise their electronic components would be subjected to irreparable damage.
The high-energy protons in the inner Van Allen belt are thought to originate from the decay of neutrons. These neutrons are produced by the interaction of energetic cosmic rays of galactic origin with the atmosphere. Some of the neutrons are ejected upward, and a fraction of them decay into protons and electrons upon passing through the region occupied by the Van Allen belts. These protons and electrons travel in spiral paths along the flux lines of the Earth's magnetic field. At intermediate latitudes, the converging flux lines cause the particles to reverse their direction, so that they move back and forth between the magnetic poles. Collisions with atoms in the thin atmosphere eventually remove the particles from the belts, but they generally survive for about 10 years. This relatively long lifetime allows particles to accumulate in the radiation belts, providing high fluxes in spite of the small magnitude of the intrinsic source.
The inner belt merges gradually with the outer belt, where a portion of the ionization is derived from the solar wind. Clear evidence of this is provided by the observation of helium ions in addition to protons in the outer belt. It has been established that helium ions account for about 10 percent of the solar wind. (Unlike the outer zone, the inner belt contains no helium ions.) The flux of electrons in the outer belt can vary over time intervals as short as a few days. These changes appear to correlate with periods of strong magnetic disturbance. They are not, however, as yet well understood.
Studies of the outer magnetosphere and related phenomena continue to draw attention. Space physicists find it a challenging task to unravel the mysteries of the ever-changing cosmic laboratory and hope to continue the saga initiated by Van Allen and other pioneers of the space age.

Michael B. McElroy
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):395–411 (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. 1–18, 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, 7:22–28 (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).