atmosphereless, adj.
/at"meuhs fear'/, n., v., atmosphered, atmosphering.
1. the gaseous envelope surrounding the earth; the air.
2. this medium at a given place.
3. Astron. the gaseous envelope surrounding a heavenly body.
4. Chem. any gaseous envelope or medium.
5. a conventional unit of pressure, the normal pressure of the air at sea level, about 14.7 pounds per square inch (101.3 kilopascals), equal to the pressure exerted by a column of mercury 29.92 in. (760 mm) high. Abbr.: atm.
6. a surrounding or pervading mood, environment, or influence: an atmosphere of impending war; a very tense atmosphere.
7. the dominant mood or emotional tone of a work of art, as of a play or novel: the chilly atmosphere of a ghost story.
8. a distinctive quality, as of a place; character: The old part of town has lots of atmosphere.
9. to give an atmosphere to: The author had cleverly atmosphered the novel for added chills.
[1630-40; < NL atmosphaera. See ATMO-, -SPHERE]

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Gaseous envelope that surrounds the Earth.

Near the surface it has a well-defined chemical composition (see air). In addition to gases, the atmosphere contains solid and liquid particles in suspension. Scientists divide the atmosphere into five main layers: in ascending order, the troposphere (surface to 6–8 mi, or 10–13 km); the stratosphere (4–11 mi, or 6–17 km, to about 30 mi, or 50 km); the mesosphere (31–50 mi, or 50–80 km); the thermosphere (50–300 mi, or 80–480 km); and the exosphere (from 300 mi and gradually dissipating). Most of the atmosphere consists of neutral atoms and molecules, but in the ionosphere a significant fraction is electrically charged. The ionosphere begins near the top of the stratosphere but is most distinct in the thermosphere. See also ozone layer.

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▪ gaseous envelope
 the gas and aerosol envelope that extends from the ocean, land, and ice-covered surface of a planet outward into space. The density of the atmosphere decreases outward, because the gravitational attraction of the planet, which pulls the gases and aerosols (microscopic suspended particles of dust, soot, smoke, or chemicals) inward, is greatest close to the surface. Atmospheres of some planetary bodies, such as Mercury, are almost nonexistent, as the primordial atmosphere has escaped the relatively low gravitational attraction of the planet and has been released into space. Other planets, such as Venus, Earth, Mars, and the giant outer planets of the solar system, have retained an atmosphere. In addition, Earth's atmosphere has been able to contain water in each of its three phases (solid, liquid, and gas), which has been essential for the development of life on the planet.

      The evolution of Earth's current atmosphere is not completely understood. It is thought that the current atmosphere resulted from a gradual release of gases both from the planet's interior and from the metabolic activities of life-forms—as opposed to the primordial atmosphere, which developed by outgassing (venting) during the original formation of the planet. Current volcanic (volcanism) gaseous emissions include water vapour (H2O), carbon dioxide (CO2), sulfur dioxide (SO2), hydrogen sulfide (H2S), carbon monoxide (CO), chlorine (Cl), fluorine (F), and diatomic nitrogen (N2; consisting of two atoms in a single molecule), as well as traces of other substances. Approximately 85 percent of volcanic emissions are in the form of water vapour. In contrast, carbon dioxide is about 10 percent of the effluent.

      During the early evolution of the atmosphere on Earth, water must have been able to exist as a liquid, since the oceans have been present for at least three billion years. Given that solar output four billion years ago was only about 60 percent of what it is today, enhanced levels of carbon dioxide and perhaps ammonia (NH3) must have been present in order to retard the loss of infrared radiation into space. The initial life-forms that evolved in this environment must have been anaerobic (i.e., surviving in the absence of oxygen). In addition, they must have been able to resist the biologically destructive ultraviolet radiation in sunlight, which was not absorbed by a layer of ozone (ozonosphere) as it is now.

      Once organisms developed the capability for photosynthesis, oxygen was produced in large quantities. The buildup of oxygen in the atmosphere also permitted the development of the ozone layer (ozonosphere) as O2 molecules were dissociated into monatomic oxygen (O; consisting of single oxygen atoms) and recombined with other O2 molecules to form triatomic ozone molecules (O3). The capability for photosynthesis arose in primitive forms of plants between two and three billion years ago. Previous to the evolution of photosynthetic organisms, oxygen was produced in limited quantities as a by-product of the decomposition of water vapour by ultraviolet radiation.

      The current molecular composition of Earth's (Earth) atmosphere is diatomic nitrogen (N2), 78.08 percent; diatomic oxygen (O2), 20.95 percent; argon (A), 0.93 percent; water (H20), about 0 to 4 percent; and carbon dioxide (CO2), 0.038 percent. Inert gases such as neon (Ne), helium (He), and krypton (Kr) and other constituents such as nitrogen oxides, compounds of sulfur, and compounds of ozone are found in lesser amounts.

      This article provides an overview of the physical forces that drive Earth's atmospheric processes, the structure of the Earth's atmosphere, and the instrumentation used to measure the Earth's atmosphere. For a full description of the processes that created the current atmosphere on Earth, see evolution of the atmosphere (atmosphere, evolution of). For information on the long-term conditions of the atmosphere as they are experienced at the surface of Earth, see climate. For a description of the highest regions of the atmosphere, where conditions are set largely by the presence of charged particles, see ionosphere and magnetosphere.

Surface budgets

Energy budget
 Earth's atmosphere is bounded at the bottom by water and land—that is, by the surface of Earth. Heating of this surface is accomplished by three physical processes—radiation (electromagnetic radiation), conduction (thermal conduction), and convection—and the temperature at the interface of the atmosphere and surface is a result of this heating.

      The relative contributions of each process depend on the wind, temperature, and moisture structure in the atmosphere immediately above the surface, the intensity of solar insolation, and the physical characteristics of the surface. The temperature occurring at this interface is of critical importance in determining how suitable a location is for different forms of life.

      The temperature of the atmosphere and surface is influenced by electromagnetic radiation, and this radiation is traditionally divided into two types: insolation from the Sun and emittance from the surface and the atmosphere. Insolation is frequently referred to as shortwave radiation; it falls primarily within the ultraviolet and visible portions of the electromagnetic spectrum and consists predominantly of wavelengths (wavelength) of 0.39 to 0.76 micrometres (0.00002 to 0.00003 inch). Radiation emitted from Earth is called longwave radiation; it falls within the infrared portion of the spectrum and has typical wavelengths of 4 to 30 micrometres (0.0002 to 0.001 inch). Wavelengths of radiation emitted by a body depend on the temperature of the body, as specified by Planck's radiation law. The Sun, with its surface temperature of around 6,000 kelvins (K; about 5,725 °C, or 10,337 °F), emits at a much shorter wavelength than does Earth, which has lower surface and atmospheric temperatures around 250 to 300 K (−23 to 27 °C, or −9.4 to 80.6 °F).

      A fraction of the incoming shortwave radiation is absorbed by atmospheric gases, including water vapour, and warms the air directly, but in the absence of clouds (cloud) most of this energy reaches the surface. The scattering of a fraction of the shortwave radiation—particularly of the shortest wavelengths by air molecules in a process called Rayleigh scattering—produces Earth's blue skies.

      When tall thick clouds are present, a large percentage (up to about 80 percent) of the insolation is reflected back into space. (The fraction of reflected shortwave radiation is called the cloud albedo.) Of the solar radiation reaching Earth's surface, some is reflected back into the atmosphere. Values of the surface albedo range as high as 0.95 for fresh snow to 0.10 for dark, organic soils. On land, this reflection occurs entirely at the surface. In water, however, albedo depends on the angle of the Sun's rays and the depth of the water column. If the Sun's rays strike the water surface at an oblique angle, albedo may be higher than 0.85; if these rays are more direct, only a small portion, perhaps as low as 0.02, is reflected, while the rest of the insolation is scattered within the water column and absorbed. Shortwave radiation penetrates a volume of water to significant depths (up to several hundred metres) before the insolation is completely attenuated. The heating by solar radiation in water is distributed through a depth, which results in smaller temperature changes at the surface of the water than would occur with the same insolation over an equal area of land.

      The amount of solar radiation reaching the surface depends on latitude, time of year, time of day, and orientation of the land surface with respect to the Sun. In the Northern Hemisphere north of 23°30′, for example, solar insolation at local noon is less on slopes facing the north than on land oriented toward the south.

      Solar radiation is made up of direct and diffuse radiation. Direct shortwave radiation reaches the surface without being absorbed or scattered from its line of propagation by the intervening atmosphere. The image of the Sun's disk as a sharp and distinct object represents that portion of the solar radiation that reaches the viewer directly. Diffuse radiation, in contrast, reaches the surface after first being scattered from its line of propagation. On an overcast day, for example, the Sun's disk is not visible, and all of the shortwave radiation is diffuse.

      Long-wave radiation is emitted by the atmosphere and propagates both upward and downward. According to the Stefan-Boltzmann law (Stefan–Boltzmann law), the total amount of long-wave energy emitted is proportional to the fourth power of the temperature of the emitting material (e.g., the ground surface or the atmospheric layer). The magnitude of this radiation reaching the surface depends on the temperature at the height of emission and the amount of absorption that takes place between the height of emission and the surface. A larger fraction of the long-wave radiation is absorbed when the intervening atmosphere holds large amounts of water vapour and carbon dioxide. Clouds with liquid water concentrations near 2.5 grams per cubic metre absorb almost 100 percent of the long-wave radiation within a depth of 12 metres (40 feet) into the cloud. Clouds with lower liquid water concentrations require greater depths before complete absorption is attained (e.g., a cloud with a water content of 0.05 gram per cubic metre requires about 600 metres [about 2,000 feet] for complete absorption). Clouds that are at least this thick emit long-wave radiation from their bases downward to Earth's surface. The amount of long-wave radiation emitted corresponds to the temperature of the lowest levels of the cloud. (Clouds with warmer bases emit more long-wave radiation downward than colder clouds.)

      The magnitude of heat flux by conduction below a surface depends on the thermal conductivity and the vertical gradient of temperature in the material beneath the surface. Soils such as dry peat, which has very low thermal conductivity (i.e., 0.06 watt per metre per K), permit little heat flux. In contrast, concrete has a thermal conductivity about 75 times as large (i.e., 4.60 watts per metre per K) and allows substantial heat flux. In water, the thermal conductivity is relatively unimportant, since, in contrast to land surfaces, insolation extends to substantial depths in the water; in addition, water can be mixed vertically.

  Vertical mixing (convection) occurs in the atmosphere as well as in bodies of water. This process of mixing is also referred to as turbulence (atmospheric turbulence). It is a mechanism of heat flux that occurs in the atmosphere in two forms. When the surface is substantially warmer than the overlying air, mixing will spontaneously occur in order to redistribute the heat. This process, referred to as free convection, occurs when the environmental lapse rate (the rate of change of an atmospheric variable, such as temperature or density, with increasing altitude) of temperature decreases at a rate greater than 1 °C per 100 metres (approximately 1 °F per 150 feet). This rate is called the adiabatic lapse rate (adiabatic process) (the rate of temperature change occurring within a rising or descending air parcel). In the ocean, the temperature increase with depth that results in free convection is dependent on the temperature, salinity, and depth of the water. For example, if the surface has a temperature of 20 °C (68 °F) and a salinity of 34.85 parts per thousand, an increase in temperature with depth of greater than about 0.19 °C per km (0.55 °F per mile) just below in the upper layers of the ocean will result in free convection. In the atmosphere, the temperature profile with height determines whether free convection occurs or not. In the ocean, free convection depends on the temperature and salinity profile with depth. Colder and more saline conditions in a surface parcel of water, for example, make it more likely for that parcel to sink spontaneously and thus become part of the process of free convection.

      Mixing can also occur because of the shear stress of the wind on the surface. Shear stress is the pulling force of a fluid moving in one direction as it passes close to a fluid or object moving in another. As a result of surface friction, the average wind velocity at Earth's surface must be zero unless that surface is itself moving, such as in rivers or ocean currents. Winds (wind) above the surface decelerate when the vertical wind shear (the change in wind velocity at differing altitudes) becomes large enough to result in vertical mixing. The process by which heat and other atmospheric properties are mixed as a result of wind shear is called forced convection. Free and forced convection are also called convective and mechanical turbulence, respectively. This convection occurs as either sensible turbulent heat flux (heat directly transported to or from a surface) or latent turbulent heat flux (latent heat) (heat used to evaporate water from a surface). When this mixing does not occur, wind speeds are weak and change little with time; plumes from power-plant stacks within this layer, for example, spread very little in the vertical and remain in close proximity to the stacks.

Water budget
 The water budget at the air-surface interface is also of crucial importance in influencing atmospheric processes. The surface gains water through precipitation (rain and snow), direct condensation, and deposition (dew and frost). On land, the precipitation is often so large that some of it infiltrates into the ground or runs off into streams, rivers, lakes, and the oceans. Some of the precipitation remaining on the surface, such as in puddles or on vegetation, immediately evaporates back into the atmosphere.

      Liquid water in the soil is also converted to water vapour by transpiration from the leaves and stems of plants and by evaporation. The roots of vegetation may extract water from within the soil and emit it through stoma, or small openings, on the leaves. In addition, water may be evaporated from the surface of the soil directly, when groundwater from below is diffused upward. Evaporation (vaporization) occurs at the surface of water bodies at a rate that is inversely proportional to the relative humidity immediately above the surface. Evaporation is rapid in dry air but much slower when the lowest levels of the atmosphere are close to saturation. Evaporation from soils is dependent on the rate at which moisture is supplied by capillary suction within the soil, whereas transpiration from vegetation is dependent on both the water available within the root zone of plants and whether the stoma are open on the leaf surfaces. Water that evaporates and transpires into the atmosphere is often transported long distances before it precipitates out.

      The input, transport, and removal of water from the atmosphere is part of the hydrologic cycle. At any one time, only a very small fraction of Earth's water is present within the atmosphere; if all the atmospheric water was condensed out, it would cover the surface of the planet only to an average of about 2.5 cm (1 inch).

Nitrogen budget (nitrogen cycle)
 The nitrogen budget involves the chemical transformation of diatomic nitrogen (N2), which makes up 78 percent of the atmospheric gases, into compounds containing ammonium (NH+), nitrite (NO2), and nitrate (NO3). In a process called nitrification, or nitrogen fixation, bacteria such as Rhizobium living within nodules on the roots of peas, clover, and other legumes (legume) convert diatomic nitrogen gas to ammonia. A small amount of nitrogen is also fixed by lightning. Ammonia may be further transformed by other bacteria into nitrites and nitrates and used by plants for growth. These compounds are eventually converted back to N2 after the plants die or are eaten by denitrifying bacteria. These bacteria, in their consumption of plants and both the excrement and corpses of plant-eating animals, convert much of the nitrogen compounds back to N2. Some of these compounds are also converted to N2 by a series of chemical processes associated with ultraviolet light from the Sun. The combustion of petroleum by motor vehicles also produces oxides of nitrogen, which enhance the natural concentrations of these compounds. smog, which occurs in many urban areas, is associated with substantially higher levels of nitrogen oxides.

Sulfur budget (sulfur cycle)
 The sulfur budget is also of major importance. Sulfur is put into the atmosphere as a result of weathering of sulfur-containing rocks and by intermittent volcanic emissions. Organic forms of sulfur are incorporated into living organisms and represent an important component in both the structure and the function of proteins. Sulfur also appears in the atmosphere as the gas sulfur dioxide (SO2) and as part of particulate compounds containing sulfate (SO4). Alone, both are directly dry-deposited or precipitated out onto Earth's surface. When wetted, these compounds are converted to caustic sulfuric acid (H2SO4).

      Since the beginning of the Industrial Revolution, human activities have injected significant quantities of sulfur into the atmosphere through the combustion of fossil fuels (fossil fuel). In and near regions of urbanization and heavy industrial activity, the enhanced deposition and precipitation of sulfur in the form of sulfuric acid, and of nitrogen oxides in the form of nitric acid (HNO3), resulting from vehicular emissions, have been associated with damage to fish populations, forests, statues, and building exteriors. The conversion of sulfur and nitrogen oxides to acids such as H2SO4 and HNO3 is commonly known as the acid rain problem. Sulfur and nitrogen oxides are precipitated in rain, snow, and dry deposition (deposition to the surface during dry weather).

Carbon budget (carbon cycle)
 The carbon budget in the atmosphere is of critical importance to climate and to life. Carbon appears in Earth's atmosphere primarily as carbon dioxide (CO2) produced naturally by the respiration of living organisms, the decay of these organisms, the weathering of carbon-containing rock strata, and volcanic emissions. Plants utilize CO2, water, and solar insolation to convert CO2 to diatomic oxygen (O2). This process, known as photosynthesis, can result in local reductions of CO2 of tens of parts per million within vegetation canopies. In contrast, nighttime respiration occurring when photosynthesis is not active can increase CO2 concentrations. These concentrations may even double within dense tropical forest canopies for short periods before sunrise. On the global scale, seasonal variations of about 1 percent occur as a result of CO2 uptake from photosynthesis, plant respiration, and soil respiration. Atmospheric CO2 is primarily absorbed in the Northern Hemisphere during the growing season (spring to autumn). CO2 is also absorbed by ocean waters; the rate of exchange to the ocean is greater for colder than for warmer waters. Currently CO2 makes up about 0.03 percent of the gaseous composition of the atmosphere.

      In the geologic past, CO2 levels have been significantly higher than they are today and have had a significant effect on both climate and ecology. During the Carboniferous Period (360 to 300 million years ago), for example, moderately warm and humid climates combined with high concentrations of CO2 were associated with extensive lush vegetation. After these plants died and decomposed, they were converted to sedimentary rocks that eventually became the coal deposits currently used for industrial combustion.

      In the atmosphere, certain wavelengths of long-wave radiation are absorbed and then reemitted by CO2. Since the lower levels of the atmosphere are warmer than layers higher up, the absorption of upward-propagating electromagnetic radiation, and a reemission of a portion of it back downward, permits the lower atmosphere to remain warmer than it would be otherwise. The association of higher concentrations of CO2 in the air with a warmer lower troposphere is commonly referred to as the greenhouse effect. (The name is inaccurate—an actual greenhouse is warmed primarily because solar radiation enters through the glass, which retains the heated air and prevents the mixing of cooler air into the greenhouse from above.) In recent years, there has been increasing concern that the release of CO2 through the burning of coal and other fossil fuels will warm the lower atmosphere, a phenomenon commonly referred to as global warming. Water vapour is a more efficient greenhouse gas than carbon dioxide. However, since H2O is ubiquitous, occurring in its three phases (solid, liquid, and gas), and since CO2 is also a biogeochemically active gas, global temperature changes are both explained and predicted by changes in the atmospheric concentration of CO2.

Vertical structure of the atmosphere
      Earth's atmosphere is segmented into two major zones. The homosphere is the lower of the two and the location in which turbulent mixing dominates the molecular diffusion of gases. In this region, which occurs below 100 km (about 60 miles) or so, the composition of the atmosphere tends to be independent of height. Above 100 km, in the zone called the heterosphere, various atmospheric gases are separated by molecular mass, with the lighter gases being concentrated in the highest layers. Above 1,000 km (about 600 miles), helium and hydrogen are the dominant species. Diatomic nitrogen (N2), a relatively heavy gas, drops off rapidly with height and exists in only trace amounts at 500 km (300 miles) and above. This decrease in the concentration of heavier gases with height is largest during periods of low Sun activity, when temperatures within the heterosphere are relatively low. The transition zone, located at a height of around 100 km between the homosphere and heterosphere, is called the turbopause.

      The atmosphere can be further divided into several distinct layers defined by changes in air temperature with increasing height. These layers are described below in order of increasing height above the surface.

 The lowest portion of the atmosphere is the troposphere, a layer where temperature generally decreases with height. This layer contains most of Earth's clouds and is the location where weather primarily occurs.

Planetary boundary layer
      The lower levels of the troposphere are usually strongly influenced by Earth's surface. This sublayer, known as the planetary boundary layer, is that region of the atmosphere in which the surface influences temperature, moisture, and wind velocity through the turbulent transfer of mass. As a result of surface friction, winds in the planetary boundary layer are usually weaker than above and tend to blow toward areas of low pressure. For this reason, the planetary boundary layer has also been called an Ekman layer, for Swedish oceanographer Vagn Walfrid Ekman (Ekman, V Walfrid), a pioneer in the study of the behaviour of wind-driven ocean currents.

      Under clear, sunny skies over land, the planetary boundary layer tends to be relatively deep as a result of the heating of the ground by the Sun and the resultant generation of convective turbulence. During the summer, the planetary boundary layer can reach heights of 1 to 1.5 km (0.6 to 1 mile) above the land surface—for example, in the humid eastern United States—and up to 5 km (3 miles) in the southwestern desert. Under these conditions, when unsaturated air rises and expands, the temperature decreases at the dry adiabatic lapse rate (9.8 °C per kilometre, or roughly 23 °F per mile) throughout most of the boundary layer. Near Earth's heated surface, air temperature decreases superadiabatically (at a lapse rate greater than the dry adiabatic lapse rate). In contrast, during clear, calm nights, turbulence tends to cease, and radiational cooling (net loss of heat) from the surface results in an air temperature that increases with height above the surface.

      When the rate of temperature decrease with height exceeds the adiabatic lapse rate for a region of the atmosphere, turbulence (atmospheric turbulence) is generated. This is due to the convective overturn (convection) of the air as the warmer lower-level air rises and mixes with the cooler air aloft. In this situation, since the environmental lapse rate is greater than the adiabatic lapse rate, an ascending parcel of air remains warmer than the surrounding ambient air even though the parcel is both cooling and expanding. Evidence of this overturn is produced in the form of bubbles, or eddies (eddy), of warmer air. The larger bubbles often have sufficient buoyant energy to penetrate the top of the boundary layer. The subsequent rapid air displacement brings air from aloft into the boundary layer, thereby deepening the layer. Under these conditions of atmospheric instability, the air aloft cools according to the environmental lapse rate faster than the rising air is cooling at the adiabatic lapse rate. The air above the boundary layer replaces the rising air and undergoes compressional warming as it descends. As a result, this entrained air heats the boundary layer.

      The ability of the convective bubbles to break through the top of the boundary layer depends on the environmental lapse rate aloft. The upward movement of penetrative bubbles will decrease rapidly if the parcel quickly becomes cooler than the ambient environment that surrounds it. In this situation, the air parcel will become less buoyant with additional ascent. The height that the boundary layer attains on a sunny day, therefore, is strongly influenced by the intensity of surface heating and the environmental lapse rate just above the boundary layer. The more rapidly a rising turbulent bubble cools above the boundary layer relative to the surrounding air, the lower the chance that subsequent turbulent bubbles will penetrate far above the boundary layer. The top of the daytime boundary layer is referred to as the mixed-layer inversion.

      On clear, calm nights, radiational cooling results in a temperature increase with height. In this situation, known as a nocturnal inversion, turbulence is suppressed by the strong thermal stratification. Thermally stable conditions occur when warmer air overlies cooler, denser air. Over flat terrain, a nearly laminar wind flow (a pattern where winds from an upper layer easily slide past winds from a lower layer) can result. The depth of the radiationally cooled layer of air depends on a variety of factors, such as the moisture content of the air, soil and vegetation characteristics, and terrain configuration. In a desert environment, for instance, the nocturnal inversion tends to be found at greater heights than in a more humid environment. The inversion in more humid environments occurs at a lower altitude because more long-wave radiation emitted by the surface is absorbed by numerous available water molecules and reemitted back toward the surface. As a result, the lower levels of the troposphere are prevented from cooling rapidly. If the air is moist and sufficient near-surface cooling occurs, water vapour will condense into what is called “radiation fog.”

Wind-generated turbulence
      During windy (wind) conditions, the mechanical production of turbulence becomes important. Turbulence eddies produced by wind shear tend to be smaller in size than the turbulence bubbles produced by the rapid convection of buoyant air. Within a few tens of metres of the surface during windy conditions, the wind speed increases dramatically with height. If the winds are sufficiently strong, the turbulence generated by wind shear can overshadow the resistance of layered, thermally stable air.

      In general, there tends to be little turbulence above the boundary layer in the troposphere. Even so, there are two notable exceptions. First, turbulence is produced near jet streams (jet stream), where large velocity shears exist both within and adjacent to cumuliform clouds. In these locations, buoyant turbulence occurs as a result of the release of latent heat. Second, pockets of buoyant turbulence may be found at and just above cloud tops. In these locations, the radiational cooling of the clouds destabilizes pockets of air and makes them more buoyant. clear-air turbulence (CAT) is frequently reported when aircraft fly near one of these regions of turbulence generation.

      The top of the troposphere, called the tropopause, corresponds to the level in which the pattern of decreasing temperature with height ceases. It is replaced by a layer that is essentially isothermal (of equal temperature). In the tropics and subtropics, the tropopause is high, often reaching to about 18 km (11 miles), as a result of vigorous vertical mixing of the lower atmosphere by thunderstorms (thunderstorm). In polar regions, where such deep atmospheric turbulence is much less frequent, the tropopause is often as low as 8 km (5 miles). Temperatures at the tropopause range from as low as −80 °C (−112 °F) in the tropics to −50 °C (−58 °F) in polar regions.

Cloud formation within the troposphere
      The region above the planetary boundary layer is commonly known as the free atmosphere. Winds at this volume are not directly retarded by surface friction. Clouds (cloud) occur most frequently in this portion of the troposphere, though fog and clouds that impinge or develop over elevated terrain often occur at lower levels.

      There are two basic types of clouds: cumuliform and stratiform. Both cloud types develop when clear air ascends, cooling adiabatically as it expands until either water begins to condense or deposition occurs. Water undergoes a change of state from gas to liquid under these conditions, because cooler air can hold less water vapour than warmer air. For example, air at 20 °C (68 °F) can contain almost four times as much water vapour as at 0 °C (32 °F) before saturation takes place and water vapour condenses into liquid droplets.

  Stratiform clouds occur as saturated air is mechanically forced upward and remains colder than the surrounding clear air at the same height. In the lower troposphere, such clouds are called stratus. advection fog is a stratus cloud with a base lying at Earth's surface. In the middle troposphere, stratiform clouds are known as altostratus. In the upper troposphere, the terms cirrostratus and cirrus are used. The cirrus cloud type refers to thin, often wispy, cirrostratus clouds. Stratiform clouds that both extend through a large fraction of the troposphere and precipitate are called nimbostratus.

      Cumuliform clouds occur when saturated air is turbulent. Such clouds, with their bubbly turreted shapes, exhibit the small-scale up-and-down behaviour of air in the turbulent planetary boundary layer. Often such clouds are seen with bases at or near the top of the boundary layer as turbulent eddies generated near Earth's surface reach high enough for condensation to occur.

      Cumuliform clouds will form in the free atmosphere if a parcel of air, upon saturation, is warmer than the surrounding ambient atmosphere. Since this air parcel is warmer than its surroundings, it will accelerate upward, creating the saturated turbulent bubble characteristic of a cumuliform cloud. Cumuliform clouds, which reach no higher than the lower troposphere, are known as cumulus humulus when they are randomly distributed and as stratocumulus when they are organized into lines. Cumulus congestus clouds extend into the middle troposphere, while deep, precipitating cumuliform clouds that extend throughout the troposphere are called cumulonimbus. Cumulonimbus (thunderstorm) clouds are also called thunderstorms, since they usually have lightning and thunder associated with them. Cumulonimbus clouds develop from cumulus humulus and cumulus congestus clouds.

Stratosphere and mesosphere
      The stratosphere is located above the troposphere and extends up to about 50 km (30 miles). Above the tropopause and the isothermal layer in the lower stratosphere, temperature increases with height. Temperatures as high as 0 °C (32 °F) are observed near the top of the stratosphere. The observed increase of temperature with height in the stratosphere results in strong thermodynamic stability with little turbulence and vertical mixing. The warm temperatures and very dry air result in an almost cloud-free volume. The infrequent clouds that do occur are called nacreous, or mother-of-pearl, clouds because of their striking iridescence, and they appear to be composed of both ice and supercooled water. These clouds form up to heights of 30 km (19 miles).

      The pattern of temperature increase with height in the stratosphere is the result of solar heating as ultraviolet radiation in the wavelength range of 0.200 to 0.242 micrometre dissociates diatomic oxygen (O2). The resultant attachment of single oxygen atoms to O2 produces ozone (O3). Natural stratospheric ozone is produced mainly in the tropical and middle latitudes. Regions of nearly complete ozone depletion, which have occurred in the Antarctic during the spring, are associated with nacreous clouds, chlorofluorocarbons (chlorofluorocarbon) (CFCs), and other pollutants from human activities. These regions are more commonly known as ozone holes. Ozone is also transported downward into the troposphere, primarily in the vicinity of the polar front.

      The stratopause caps the top of the stratosphere, separating it from the mesosphere near 45–50 km (28–31 miles) in altitude and a pressure of 1 millibar (approximately equal to 0.75 mm of mercury at 0 °C, or 0.03 inch of mercury at 32 °F). In the mesosphere, temperatures again decrease with increasing altitude. Unlike the situation in the stratosphere, vertical air currents in the mesosphere are not strongly inhibited. Ice crystal clouds, called noctilucent clouds (noctilucent cloud), occasionally form in the upper mesosphere. Above the mesopause, a region occurring at altitudes near 85 to 90 km (50 to 55 miles), temperature again increases with height in a layer called the thermosphere.

      Temperatures in the thermosphere range from near 500 K (approximately 227 °C, or 440 °F) during periods of low sunspot activity to 2,000 K (1,725 °C, or 3,137 °F) when the Sun is active. The thermopause, defined as the level of transition to a more or less isothermal temperature profile at the top of the thermosphere, occurs at heights of around 250 km (150 miles) during quiet Sun periods and almost 500 km (300 miles) when the Sun is active. Above 500 km, molecular collisions are infrequent enough that temperature is difficult to define.

      The portion of the thermosphere where charged particles (ions (ion)) are abundant is called the ionosphere (ionosphere and magnetosphere). These ions result from the removal of electrons from atmospheric gases by solar ultraviolet radiation. Extending from about 80 to 300 km (about 50 to 185 miles) in altitude, the ionosphere is an electrically conducting region capable of reflecting radio signals back to Earth.

      Maximum ion density, a condition that makes for efficient radio transmission, occurs within two sublayers: the lower E region, which exists from 90 to 120 km (about 55 to 75 miles) in altitude; and the F region, which exists from 150 to 300 km (about 90 to 185 miles) in altitude. The F region has two maxima (i.e., two periods of highest ion density) during daylight hours, called F1 and F2. Both the F1 and F2 regions possess high ion density and are strongly influenced by both solar activity and time of day. Of these, the F2 region is the more variable of the two and may reach an ion density as high as 106 electrons per cubic centimetre. Shortwave radio transmissions, capable of reaching around the world, take advantage of the ability of layers in the ionosphere to reflect certain wavelengths of electromagnetic radiation. In addition, electrical discharges from the tops of thunderstorms into the ionosphere, called transient luminous events, have been observed.

magnetosphere and exosphere
 Above approximately 500 km (300 miles), the motion of ions (ion) is strongly constrained by the presence of Earth's magnetic field. This region of Earth's atmosphere, called the magnetosphere, is compressed by the solar wind on the daylight side of the planet and stretched outward in a long tail on the night side. The colourful auroral (aurora) displays often seen in polar latitudes are associated with bursts of high-energy particles generated by the Sun. When these particles are influenced by the magnetosphere, some are subsequently injected into the lower ionosphere.

      The layer above 500 km is referred to as the exosphere, a region in which at least half of the upward-moving molecules do not collide with one another. In contrast, these molecules follow long ballistic trajectories and may exit the atmosphere completely if their escape velocities are high enough. The loss rate of molecules through the exosphere is critical in determining whether Earth or any other planetary body retains an atmosphere.

Horizontal structure of the atmosphere

Distribution of heat from the Sun
      The primary driving force for the horizontal structure of Earth's atmosphere is the amount and distribution of solar radiation that comes in contact with the planet. Earth's orbit around the Sun is an ellipse, with a perihelion (closest approach) of 147.5 million km (91.7 million miles) in early January and an aphelion (farthest distance) of 152.6 million km (94.8 million miles) in early July. As a result of Earth's elliptical orbit, the time between the autumnal equinox and the following vernal equinox (about September 22 to about March 21) is almost one week shorter than the remainder of the year in the Northern Hemisphere. This results in a shorter astronomical winter in the Northern Hemisphere than in the Southern Hemisphere.

      Earth rotates once every 24 hours around an axis that is tilted at an angle of 23°30′ with respect to the plane of its orbit around the Sun. As a result of this tilt, during the summer season of either the Northern or the Southern Hemisphere, the Sun's rays are more direct at a given latitude than they are during the winter season. Poleward of latitudes 66°30′ N and 66°30′ S, the tilt of the planet is such that for at least one complete day (at 66°30′) and as long as six months (at 90°), the Sun is above the horizon during the summer season and below the horizon during the winter.

      As a result of this asymmetric distribution of solar heating, during the winter season the troposphere in the high latitudes becomes very cold. In contrast, during the summer at high latitudes, the troposphere warms significantly as a result of the long hours of daylight; however, owing to the oblique angle of the sunlight near the poles, the temperatures there remain relatively cool compared with middle latitudes. Equatorward of latitudes 30° N and 30° S or so, substantial radiant heating from the Sun occurs during both winter and summer seasons. The tropical troposphere, therefore, has comparatively little variation in temperature during the year.

convection, circulation, and deflection of air
      The region of greatest solar heating at the surface in the humid tropics corresponds to areas of deep cumulonimbus convection. Cumulonimbus clouds routinely form in the tropics where rising parcels of air are warmer than the surrounding ambient atmosphere. They transport water vapour, sensible heat, and Earth's rotational momentum to the upper portion of the troposphere. As a result of the vigorous convective mixing of the atmosphere, the tropopause in the lower latitudes is often very high, located some 17 to 18 km (10.5 to 11 miles) above the surface.

      Since motion upward into the stratosphere is inhibited by very stable thermal layering, the air transported upward by convection diverges toward the poles in the upper troposphere. (This divergence aloft results in a wide strip of low atmospheric pressure at the surface in the tropics, occurring in an area called the equatorial trough). As the diverted air in the troposphere moves toward the poles, it tends to retain the angular momentum of the near-equatorial region, which is large as a result of Earth's rotation. As a result, the poleward-moving air is deflected toward the right in the Northern Hemisphere and toward the left in the Southern Hemisphere.

 Upon reaching around 30° of latitude poleward of its region of origin, the upper-level air is traveling primarily toward the poles and is tending toward the east. Since motion upward is constrained by the stratosphere, the slowly cooling air must descend. The compressional warming that occurs as the air descends creates vast regions of subtropical high pressure. These regions are centred over the oceans and are characterized by strong thermodynamic stability. The sparse precipitation in these regions, a result of stability and subsidence (convergence and divergence), is associated with such great arid regions of the world as the Sahara, Atacama (Atacama Desert), Kalahari, and Sonoran (Sonoran Desert) deserts. The accumulation of air as a result of the convergence in the upper troposphere causes deep high-pressure systems, known as subtropical ridges (subtropical high), to form in these regions. Locally, these ridges are given such names as the Bermuda High, the Azores high, and the North Pacific High.

 The descending air referred to above, upon reaching the lower troposphere, is forced to diverge by the presence of Earth's surface. Some air moves poleward, while the remainder moves equatorward. In either direction, the air is deflected to the right in the Northern Hemisphere and to the left in the Southern Hemisphere. Deflection occurs because, in accordance with Newton's first law of motion (Newton's laws of motion), a parcel moving in a certain direction will retain the same motion unless acted on by an exterior force. With respect to a rotating Earth, a moving parcel conserving its momentum (i.e., not acted on by an exterior force) will appear to be deflected with respect to fixed points on the rotating Earth. As seen from a fixed point in space, such a parcel would be moving in a straight line. This apparent force on the motion of a fluid (in this case, air) is called the Coriolis effect (Coriolis force). As a result of the Coriolis effect, air tends to rotate counterclockwise around large-scale low-pressure systems and clockwise around large-scale high-pressure systems in the Northern Hemisphere. In the Southern Hemisphere, the flow direction is reversed.

      In the equatorward-moving flow, this deflection results in northeast winds north of 0° latitude and southeast winds south of that latitude. These low-level winds have been called the trade winds (trade wind) since 17th-century sailing vessels used them to travel to the Americas. The convergence region for lower-level northeast and southeast trade winds is called the intertropical convergence zone (ITCZ). The ITCZ corresponds to the equatorial trough and is the mechanism that helps generate the deep cumulonimbus clouds through convection. Cumulonimbus clouds are the main conduit transporting tropical heating into the upper troposphere.

      The circulation pattern described above—ascent in the equatorial trough, poleward movement in the upper troposphere, descent in the subtropical ridges, and equatorward movement in the trade winds—is in effect a direct heat engine, which meteorologists call the Hadley cell. This persistent circulation mechanism transports heat from the latitudes of greatest solar insolation to the latitudes of the subtropical ridges. The geographic location of the Hadley circulation moves north and south with the seasons; however, the equatorial trough lags behind for about two months owing to the thermal inertia of Earth's surface. (For a given location on Earth's surface, the highest daily temperatures are achieved just after the period of greatest insolation, since time is required to heat the ocean surface waters and the soil.)

Extratropical cyclones (extratropical cyclone)
      Poleward of the subtropical ridges, winds in the lower troposphere tend to be southwesterly in the Northern Hemisphere and northwesterly in the Southern Hemisphere, again owing to the Coriolis effect. Since warm air is being moved poleward at low altitudes, the wind flow is no longer associated with the direct heat engine of the Hadley cell. Instead, the continued transport of heat from the equatorial trough toward the poles is facilitated by large low-pressure eddies called extratropical cyclones. These phenomena develop along the polar front, which separates colder polar air from warmer tropical air, when sufficiently large temperature differences occur across the frontal boundary in the lower troposphere. The intensity of this temperature gradient is referred to as the baroclinicity of the front.

      Extratropical cyclones have three stages of expansion: the developing stage, in which an undulating wave develops along the front; the mature stage, in which sinking cold air sweeps equatorward west of the surface low-pressure centre and ascending warm air moves poleward east of the cyclone; and the occluded stage, in which the warm air is entrained within and moved above the polar air and becomes separated from the source region of the tropical air. Cyclones that progress no farther than the developing stage are referred to as wave cyclones, while extratropical lows that reach the mature and occluded stages are called baroclinically unstable waves. Extratropical storm development is referred to as cyclogenesis. Rapid extratropical cyclone development, called explosive cyclogenesis, is often associated with major winter storms and occurs when surface pressure falls by more than about 24 millibars per day. Theoretical analysis has shown that the occurrence of baroclinically unstable waves is directly proportional to the magnitude of the temperature gradient, with maximum growth for wavelengths of 3,000 to 5,000 km (1,865 to 3,100 miles). Wavelengths that are shorter are damped by horizontal mixing. The 3,000 to 5,000 km wavelength is the typical separation between high- and low- pressure synoptic weather systems in the middle and higher latitudes.

Polar fronts and the jet stream
      In the troposphere, the demarcation between polar air and warmer tropical atmosphere is usually defined by the polar front. On the poleward side of the front, the air is cold and more dense; equatorward of the front, the air is warmer and more buoyant. During the winter season, the polar front is generally located at lower latitudes and is more pronounced than in the summer.

      Cold fronts (cold front) occur at the leading edge of equatorward-moving polar air. In contrast, warm fronts are well defined at the equatorward surface position of polar air as it retreats on the eastern sides of extratropical cyclones. Equatorward-moving air behind a cold front occurs in pools of dense high pressure known as polar highs and arctic highs. The term arctic high is used to define air that originates even deeper within the high latitudes than polar highs.

      When polar air neither retreats nor advances, the polar front is called a stationary front. In the occluded stage of the life cycle of an extratropical cyclone, when cold air west of the surface low-pressure centre advances more rapidly toward the east than cold air ahead of the warm front, warmer, less-dense air is forced aloft. This frontal intersection is called an occluded front. Without exception, fronts of all types follow the movement of colder air.

      Clouds and often precipitation occur on the poleward sides of both warm and stationary fronts and whenever tropical air reaching the latitude of the polar front is forced upward over the colder air near the surface. Such fronts are defined as active fronts. Rain and snowfall from active fronts form a major part of the precipitation received in the middle and high latitudes. Precipitation in these areas occurs primarily during the winter months.

      The position of the polar front slopes upward toward colder air. This occurs because cold air tends to undercut the warmer air of tropical origin. Since cold air is more dense, atmospheric pressure decreases more rapidly with height on the poleward side of the polar front than on the warmer tropical side. This creates a large horizontal temperature contrast, which is essentially a large pressure gradient, between the polar and tropical air. In the middle and upper parts of the troposphere, this pressure gradient is responsible for the strong westerly winds occurring there. Winds created aloft circulate around a large region of upper-level low pressure near each of the poles. The centre of each low pressure region is a persistent cyclone known as the circumpolar vortex.

      The region of strongest winds, which occurs at the juncture of the tropical and polar air masses, is called the jet stream. Since the temperature contrast between the tropics and the high latitudes is greatest in the winter, the jet stream is stronger during that season. In addition, since the mid-latitudes also become colder during the winter, while tropical temperatures remain relatively unchanged, the westerly jet stream approaches latitudes of 30° during the colder season. During the warmer season in both hemispheres, the jet stream moves poleward and is located between latitudes of 50° and 60°.

      The jet stream reaches its greatest velocity at the tropopause. Above that level, a reversal of the horizontal temperature gradient occurs, which produces a reduction in the wind speeds of the jet stream at high latitudes. This causes a weakening of the westerlies with increasing height. At intervals ranging from 20 to 40 months, with a mean value of 26 months, westerly winds in the stratosphere reverse direction over low latitudes, so that an easterly flow develops. This feature is called the quasi-biennial oscillation (QBO). In addition, a phenomenon called sudden stratospheric warming, apparently the result of strong downward air motion, also occurs in the late winter and spring at high latitudes. Sudden stratospheric warming can significantly alter temperature-dependent chemical reactions of ozone and other reactive gases in the stratosphere and affect the development of such features as “ozone holes.”

 A major focus of weather forecasting in the middle and high latitudes is to forecast the movement and development of extratropical cyclones (extratropical cyclone), polar and arctic highs, and the location and intensity of subtropical ridges. Spring and fall frosts, for example, are associated with the equatorward movement of polar highs behind a cold front, while droughts and heat waves in the summer are associated with unusually strong subtropical ridges.

Effect of continents on air movement
      Preferred geographic locations exist for subtropical ridges and for the development, movement, and decay of extratropical cyclones. During the winter months in middle and high latitudes, the lower parts of the troposphere over continents often serve as reservoirs of cold air as heat is radiated into space throughout the long nights. In contrast, the oceans lose heat less rapidly, because of the large heat capacity of water, their ability to overturn as the surfaces cool and become negatively buoyant, and the movement of ocean currents such as the Gulf Stream and the Kuroshio current. Warm currents transport heat from lower latitudes poleward and tend to occur on the western sides of oceans. The lower troposphere over these warmer oceanic areas tends to be a region of relative low pressure. As a result of this juxtaposition of cold air and warm air, the eastern sides of continents and the western fringes of oceans in middle and high latitudes are the preferred locations for extratropical storm development. Over Asia in particular, the cold high-pressure system is sufficiently permanent that a persistent offshore flow called the winter monsoon occurs.

      An inverse type of flow develops in the summer as the continents heat more rapidly than their adjacent oceanic areas. Continental areas tend to become regions of relative low pressure, while high pressure in the lower troposphere becomes more prevalent offshore. As the winds travel from areas of higher pressure to areas of lower pressure, a persistent onshore flow develops over large landmasses in the lower troposphere. The result of this heating is referred to as the summer monsoon. The leading edge of this monsoon is associated with a feature called the monsoon trough, a region of low atmospheric pressure at sea level. Tropical moisture carried onshore by the summer monsoon often results in copious rainfall. The village of Cherrapunji in northeastern India, for instance, recorded over 9 metres (about 30 feet) of rain in one month (July 1861) owing to the Indian summer monsoon.

      As a result of the continental effect, the subtropical ridge is segmented into surface high-pressure cells. In the summer, large landmasses in the subtropics tend to be centres of relative low pressure as a result of strong solar heating. As a consequence, persistent high-pressure cells, such as the Bermuda and Azores highs (Azores high), occur over the oceans. The oval shape of these high-pressure cells creates a thermal structure on their eastern sides that differs from the thermal structure on their western sides in the lower troposphere. On the eastern side, subsidence from the Hadley circulation is enhanced by the tendency of air to preserve its angular momentum on the rotating Earth. Owing to the enhanced descent of air over the eastern parts of the oceans, landmasses adjacent to these areas (typically the western sides of continents) tend to be deserts, such as those found in northwestern and southwestern Africa and along western coastal Mexico.

Effect of oceans (ocean) on air movement
      The arid conditions found along the western coasts of continents in subtropical latitudes are further enhanced by the influence of the equatorward surface air flow on the ocean currents. This flow exerts a shearing stress on the ocean surface, which results in the deflection of the upper layer of water above the thermocline to the right in the Northern Hemisphere and to the left in the Southern Hemisphere. (This deflection is also the result of the Coriolis effect; water from both hemispheres moves westward when displaced toward the Equator.) As warmer surface waters are carried away by this offshore ocean airflow, cold water from below the thermocline rises to the surface in a process called upwelling. Upwelling creates areas of cold coastal surface waters that stabilize the lower troposphere and reduce the chances for convection. Lower convection in turn reduces the likelihood for precipitation, although fogs and low stratus clouds are common. Upwelling regions are also associated with enriched sea life, as oxygen and organic nutrients are transported upward from the depths toward the surface of the ocean.

      During periods when the intertropical convergence zone (ITCZ) is located near the Equator, trade winds from the northeast and southeast converge there. The westward-moving winds cause the displacement of surface ocean waters away from the Equator such that the deeper, colder waters move to the surface. In the central and eastern Pacific Ocean near the Equator, when this upwelling is stronger than average, the event is called La Niña. When the trade winds weaken in this region, however, warmer-than-average surface conditions occur, and upwelling is weaker than usual. This event is called El Niño. Changes in ocean surface temperatures caused by El Niño significantly affect where cumulonimbus clouds form in the ITCZ and, therefore, the geographic structure of the Hadley cell. During periods when El Niño is active, weather patterns across the entire Earth are substantially altered.

Mountain barriers
      North-south-oriented mountain barriers, such as the Rockies (Rocky Mountains) and the Andes (Andes Mountains), and large massifs, such as the Plateau of Tibet (Tibet, Plateau of), also influence atmospheric flow. When the general westerly flow in the mid-latitudes reaches these barriers, air tends to be blocked. It is transported poleward west of the terrain and toward the Equator east of the obstacle. Air forced up the slopes of mountain barriers is often sufficiently moist to produce considerable precipitation on windward sides of mountains, whereas subsiding air on the lee slopes produces more-arid conditions. Essentially, the elevated terrain affects the atmosphere as if it were an anticyclone, a centre of high pressure. In addition, mountains prevent cold air from the continental interior from moving westward of the terrain. As a result, relatively mild weather occurs along the western coasts of continents with north-south mountain ranges when compared with continental interiors. For example, the West Coast of North America experiences milder winter weather than the Great Plains and Midwest, both of which occur at similar latitudes. In contrast, east-west mountain barriers, such as the Alps in Europe, offer little impediment to the general westerly flow of air. In these situations, milder maritime conditions extend much farther inland.

Cloud processes

 The formation of cloud droplets and cloud ice crystals is associated with suspended aerosols (aerosol), which are produced by natural processes as well as human activities and are ubiquitous in Earth's atmosphere. In the absence of such aerosols, the spontaneous conversion of water vapour into liquid water or ice crystals requires conditions with relative humidities much greater than 100 percent, with respect to a flat surface of H2O. The development of clouds in such a fashion, which occurs only in a controlled laboratory environment, is referred to as homogeneous nucleation. Air containing water vapour with a relative humidity greater than 100 percent, with respect to a flat surface, is referred to as being supersaturated. In the atmosphere, aerosols serve as initiation sites for the condensation or deposition of water vapour. Since their surfaces are of discrete sizes, aerosols reduce the amount of supersaturation required for water vapour to change its phase and are referred to as cloud condensation nuclei (condensation nucleus).

      The larger the aerosol and the greater its solubility, the lower the supersaturation percentage required for the aerosol to serve as a condensation surface. Condensation nuclei in the atmosphere become effective at supersaturations of around 0.1 to 1 percent (that is, levels of water vapour around 0.1 to 1 percent above the point of saturation). The concentration of cloud condensation nuclei in the lower troposphere at a supersaturation of 1 percent ranges from around 100 per cubic centimetre (approximately 1,600 per cubic inch) in size in oceanic air to 500 per cubic centimetre (8,000 per cubic inch) in the atmosphere over a continent. Higher concentrations occur in polluted air.

      Aerosols that are effective for the conversion of water vapour to ice crystals are referred to as ice nuclei. In contrast to cloud condensation nuclei, the most effective ice nuclei are hydrophobic (having a low affinity for water) with molecular spacings and a crystallographic structure close to that of ice.

      While cloud condensation nuclei are always readily available in the atmosphere, ice nuclei are often scarce. As a result, liquid water cooled below 0 °C (32 °F) can often remain liquid at subfreezing temperatures because of the absence of effective ice nuclei. Liquid water at temperatures less than 0 °C is referred to as supercooled water. Except for true ice crystals, which are effective at 0 °C, all other ice nuclei become effective at temperatures below freezing. In the absence of any ice nuclei, the freezing of supercooled water droplets of a few micrometres in radius, in a process called homogeneous ice nucleation, requires temperatures at or lower than −39 °C (−38 °F). While a raindrop will freeze near 0 °C, small cloud droplets have too few molecules to create an ice crystal by random chance until the molecular motion is slowed as the temperature approaches −39 °C. When ice nuclei are present, heterogeneous ice nucleation can occur at warmer temperatures.

      Ice nuclei are of three types: deposition nuclei, contact nuclei, and freezing nuclei (freezing nucleus). Deposition nuclei are analogous to condensation nuclei in that water vapour directly deposits as ice crystals on the aerosol. Contact and freezing nuclei, in contrast, are associated with the conversion of supercooled water to ice. A contact nucleus converts liquid water to ice by touching a supercooled water droplet. Freezing nuclei are absorbed into the liquid water and convert the supercooled water to ice from the inside out.

      Examples of cloud condensation nuclei include sodium chloride (NaCl) and ammonium sulfate ([NH4]2 SO2), whereas the clay mineral kaolinite is an example of an ice nuclei. In addition, naturally occurring bacteria found in decayed leaf litter can serve as ice nuclei at temperatures of less than about −4 °C (24.8 °F). In a process called cloud seeding, silver iodide, with effective ice-nucleating temperatures of less than −4 °C, has been used for years in attempts to convert supercooled water to ice crystals in regions with a scarcity of natural ice nuclei.

Liquid droplets
      The evolution of clouds that follows the formation of liquid cloud droplets or ice crystals depends on which phase of water occurs. A cloud in which only liquid water occurs (even at temperatures less than 0 °C) is referred to as a warm cloud, and the precipitation that results is said to be due to warm-cloud processes. In such a cloud, the growth of a liquid water droplet to a raindrop begins with condensation, as additional water vapour condenses in a supersaturated atmosphere. This process continues until the droplet has attained a radius of about 10 micrometres (0.0004 inch). Above this size, since the mass of the droplet increases according to the cube of its radius, further increases by condensational growth are very slow. Subsequent growth, therefore, occurs only when the cloud droplets develop at slightly different rates. Differences in growth rates have been attributed to differences in spatial variations of the initial aerosol sizes, in solubilities, and in magnitudes of supersaturation. Cloud droplets of different sizes will fall at different velocities and will collide with droplets of different radii. If the collision is hard enough to overcome the surface tension between the two colliding droplets, coalescence will occur and result in a new and larger single droplet.

      This process of cloud-droplet growth is referred to as collision-coalescence. Warm-cloud rain results when the droplets attain a sufficient size to fall to the ground. Such a raindrop (perhaps about 1 mm [0.04 inch] in radius) contains perhaps one million 10-micrometre cloud droplets. The typical radii of raindrops resulting from this type of precipitation process range up to several millimetres and have fall velocities of around 3 to 4 metres (10 to 13 feet) per second. This type of precipitation is very common from shallow cumulus clouds over tropical oceans. In these locations, the concentration of cloud condensation nuclei is so small that there is only limited competition for the available water vapour.

Precipitation of ice
      A cloud that contains ice crystals is referred to as a cold cloud, and the resulting precipitation is said to be the product of cold-cloud processes. Traditionally, this process has also been referred to as the Bergeron-Findeisen mechanism, for Swedish meteorologists Tor Bergeron (Bergeron, Tor Harold Percival) and Walter Findeisen, who introduced it in the 1930s. In this type of cloud, ice crystals can grow directly from the deposition of water vapour. This water vapour may be supersaturated with respect to ice, or it may be the result of evaporation of supercooled water and subsequent deposition onto an ice crystal. Since the saturation vapour pressure of liquid water is always greater than or equal to the saturation vapour pressure of ice, ice crystals will grow at the expense of the liquid water. For example, saturated air with respect to liquid water becomes supersaturated with respect to ice by 10 percent at −10 °C (14 °F) and by 21 percent at −20 °C (-4 °F). This results in a rapid conversion of liquid water to ice. This substantial and rapid change of phase permits large ice crystals in a cloud surrounded by a large number of supercooled cloud droplets to grow quickly (often in less than 15 minutes) from tiny ice crystals to snowflakes. These snowflakes are large enough to fall by depositional growth alone. Fall velocities of snow range up to about 2 metres per second (6.5 feet per second). Ice crystals that grow by deposition have much lower densities than solid ice because of the air pockets occurring within the volume of the crystal. This lower density differentiates snow from ice. Clouds that are completely converted to ice crystals are referred to as glaciated clouds.

      The specific form the ice crystals take depends on the temperature and the degree of supersaturation with respect to ice. At −14 °C (7 °F) and a relatively large supersaturation with respect to liquid water, for example, ice crystals with dendritic (treelike branching) patterns form. This type of ice crystal, the one usually used to represent snowflakes in photographs and drawings, experiences growth at the end of radial arms on one or more planes of the crystal. At −40 °C (−40 °F) and a supersaturation with respect to liquid water of close to 0 percent, hollow ice columns form.

      Ice crystals can also grow large enough to precipitate either by aggregation or by riming. Aggregation occurs when the arms of the ice crystals interlock and form a clump. This collection of intermingled ice crystals can occasionally reach several centimetres in diameter. Ice crystals can also grow when supercooled water freezes directly onto the crystal to form rime. With greater accumulation of dense ice on the crystal, its fall velocity increases. When the riming is substantial enough, the crystal form of the snowflake is lost and replaced by a more or less spherical particle called graupel. Smaller-sized graupels are generally referred to as snow grains. In cumulonimbus clouds during conditions where graupels are repeatedly wetted and then injected back toward high altitudes by strong updrafts, very large graupels called hail result. Hail has been observed on the ground at sizes larger than grapefruits.

      Frozen precipitation, falling to levels of the atmosphere that are much warmer than 0 °C, often melts and reaches the ground as rain. Such cold-cloud rain at the ground is usually distinguished from warm-cloud rain by its larger size. Melted hailstones, in particular, make a large-radius impact when they strike the ground. Cold-cloud rain occasionally will refreeze if a layer of subfreezing air exists near Earth's surface. When this freezing occurs in the free atmosphere, the frozen raindrops are referred to as sleet or ice pellets. When this freezing occurs only upon the impact of the raindrop with the ground, the precipitation is known as freezing rain. During ice storms, freezing rain can produce accumulations heavy enough to snap large trees and electrical lines.

lightning and optical phenomena
 The repeated collision of ice crystals and graupel in clouds is associated with the buildup of electrical charge. This electrification is particularly large in cumulonimbus clouds as a result of vigorous vertical mixing and collisions. On average, positive charges accumulate in the upper regions, while negative charges are concentrated lower down. In response to the negative charge near the cloud base, and as negatively charged rain falls toward the ground, a pocket of positive charge develops on the ground. When the difference in electric potential between positive and negative charges becomes large enough, a sudden electrical discharge (lightning) will occur. Lightning can occur between different regions of the cloud, as in intracloud lightning (thunderstorm), and between the cloud and the positively charged ground, as in cloud-to-ground lightning (thunderstorm). The passage of the lightning through the air heats it to above 30,000 K (29,725 °C, or 53,540 °F), causing a large increase in pressure. This produces a powerful shock wave that is heard as thunder.

      Sunlight that propagates through clouds and precipitation often produces fascinating optical images. Rainbows (rainbow) are produced when sunlight is diffracted into its component colours by water droplets. In addition, halos (halo) are produced by the refraction and reflection of sunlight or moonlight by ice crystals, while coronas are formed when sunlight or moonlight passes through water droplets.

Cloud research
      The presence of cloud condensation and ice nuclei in air parcels is tested by using cloud chambers in which controlled temperatures and relative humidities are specified. In the upper troposphere and lower stratosphere, aircraft fly through clouds collecting droplets and ice on collection plates or photographing their presence in the airstream. In the past, identification of the different sizes of droplets and of the various types of ice crystals was performed by a researcher in a tedious and subjective procedure. Today this analysis can be automated by computerized image assessment. On the ground, rainfall impaction molds and snow crystal impressions are made. Hailstones are also collected, since an analysis of their structure often helps define the ambient environment in which they formed. Chemical analyses of the cloud droplets, ice crystals, and precipitation are also frequently performed in order to identify natural and man-made pollutants within the different forms of water.

Measurement systems
 Methods to monitor the atmosphere are of two types—in situ measurements and remote sensing observations. In situ measurements require that the instrumentation be located directly at the point of interest and in contact with the subject of interest. In contrast, remote sensors are located some distance away from the subject of interest. Remote sensors include passive systems (instruments that receive information naturally emitted from a region in the atmosphere) and active systems (instruments that emit either acoustic or electromagnetic energy and record the characteristics of this energy after it reflects off an object or surface and returns back to the sensor).

      Within the planetary boundary layer, in situ instrumentation includes towers, tethered balloons, and surface data collection platforms. A wide range of meteorological measurements are made from this equipment, including temperature, dew-point temperature, pressure, wind velocity, long-wave and shortwave radiative fluxes, and air chemistry. Active remote-sensing observations are made, using Doppler (Doppler effect) and non-Doppler radars (radar), lidars (a type of laser that measures backscattered light), and acoustic sounders. Radars measure the backscattering of electromagnetic microwave radiation with wavelengths on the order of 3 to 10 cm (1 to 4 inches). The non-Doppler radars provide estimates of precipitation intensity, while Doppler radars can also provide estimates of wind speed and direction by detecting a shift in the frequency of an echo produced by a moving target. Shorter-wavelength Doppler radars are often able to measure winds even in clear air. Carbon dioxide lidars provide estimates of wind structure and turbulence within a few tens of kilometres of the instrument. Acoustic sounders are used primarily to monitor boundary layer depth and structure, using echo-return characteristics. Passive instrumentation includes the pyranometer, which measures direct and diffuse solar radiation, and the pyrheliometer, which samples only direct radiation from the Sun.

      Above the boundary layer, but within the troposphere, the primary standard observation platform is the radiosonde. Tethered to helium balloons, radiosondes are released twice daily (simultaneously at 0000 hours and 1200 hours Greenwich Mean Time) around the world. As a result of their use, a long-period data archive of the status of the atmosphere has been achieved. Meteorological observations from radiosondes are also applied to benchmark the numerical weather prediction models used to forecast day-to-day weather. Radiosondes measure temperature, dew-point temperature, and pressure. The position of the radiosonde can be monitored by radar tracking so that wind speed and direction as a function of height are routinely available—for this reason radiosondes are also referred to as rawinsondes. Since the 1990s, the global positioning system ( GPS) has been used to track the balloons and calculate wind speed and direction. The radiosondes are designed to have a rise rate of about 200 metres (650 feet) per minute.

      Remote-sensing systems called profilers have been developed to provide almost continuous measurements of wind and, somewhat less accurately, of moisture and temperature throughout the lowest 10 km (6 miles) of the atmosphere. Winds are estimated by using an upward-looking Doppler radar, while temperature and moisture profiles are evaluated by using a vertically pointing radiometer that measures electromagnetic emissions of selected wavelengths at various heights in the troposphere. Used in conjunction with Earth-orbiting satellite-based passive temperature and moisture radiometric soundings, as well as active lidar wind measurements, profilers complement the data collected from radiosondes.

      Aircraft also provide detailed information concerning the structure of the atmosphere. Airplanes used in field experiments, such as the Lockheed P-3 aircraft employed by the National Oceanic and Atmospheric Administration (NOAA) in the United States, are heavily instrumented and often carry Doppler radar, turbulence sensors, and in situ measurement devices for cloud water, cloud ice content, and structure. The NOAA P-3 has been used to fly through hurricanes (tropical cyclone) and other types of deep precipitating cloud systems. Commercial aircraft are used routinely to collect atmospheric data temperature and wind data. This information is communicated to weather forecasters and used in the preparation of weather map analyses.

      Lightning occurrences are monitored by using ground-based detectors. Such systems measure time, location, flash polarity, and stroke count of lightning strikes. When the observations from systems at different locations are combined, distribution maps of lightning strikes, and hence thunderstorm occurrences, can be made.

      Above the routine maximum height of the radiosonde data (above levels where atmospheric pressure drops below 100 millibars, at about 17 km [10.5 miles]), rocketsondes, rocket-borne grenades, and falling sphere experiments have been used to monitor the thermal structure of the upper atmosphere. Since these measurements occur much less frequently than radiosonde observations, however, less is known about the meteorology above the tropopause. Satellite radiometric soundings have also been used to provide temperature readings in layers in the atmosphere from near the surface up to about 25 km (16 miles) or so, although these measurements offer less vertical and spatial resolution than in situ measurements. Similarly, ground-based radar and lidars have been used to measure atmospheric characteristics in the upper atmosphere.

The atmospheres of other planets (planet)
 Astronomical bodies retain an atmosphere when their escape velocity is significantly larger than the average molecular velocity of the gases present in the atmosphere. There are 8 planets and over 160 moons in the solar system. Of these, the planets Venus, Earth, Mars, Jupiter, Saturn, Uranus, and Neptune have significant atmospheres. Pluto (a dwarf planet) may have an appreciable atmosphere, but perhaps only when its highly elliptical orbit is closest to the Sun. Of the moons, only Titan, a moon of Saturn, is known to have a thick atmosphere. Much of what is known of these planets and their moons has resulted from the Pioneer, Viking, Mariner, Voyager, and Venera space probes.

      The atmosphere of Venus is about 96 percent carbon dioxide, with surface temperatures around 737 K (464 °C, or 867 °F). Clouds on Venus are made of sulfuric acid (H2SO4) and move in an easterly circulation of about 100 metres per second (224 miles per hour). Venus itself rotates only once every 243 Earth days. Surface pressures on Venus are around 95,000 millibars. (By contrast, Earth has a sea-level pressure of around 1,000 millibars.)

       Mars, in contrast, has a thin atmosphere composed of about 95 percent carbon dioxide, with the remainder being mostly diatomic nitrogen. Traces of water vapour also occur. Mars has a mean surface air temperature estimated at 210 K (−63 °C, or −82 °F), and surface pressures hover near 6 millibars. Both water and carbon dioxide clouds are observed on Mars, and it has well-defined seasons. In addition to periodic regional and global dust storms, cyclonic storms and clouds, associated with the boundary between cold air (from the polar cap) and warm air (from the mid-latitudes), have been observed on the planet. The rotation rate of Mars is close to the rotation rate of Earth. Evidence of river channels on Mars indicates that liquid water was present and atmospheric density was much higher in the planet's geologic past.

      Along with Earth, Venus and Mars have atmospheres that were primarily formed as a result of volcanic gas emissions, although the evolution of these gases on each planet has been very different. On Mars, for example, temperatures are currently so low that most of the water vapour emitted by volcanoes has apparently been deposited as ice within the crustal soils. The closer proximity of Venus to the Sun, and the resultant higher temperatures, may have led to the loss of most of the water from that planet—most likely through the dissolution of water into hydrogen and oxygen. Hydrogen gas was lost to space; oxygen was combined with other elements through oxidation (oxidation–reduction reaction); and carbon dioxide (produced by volcanic emissions) accumulated to high concentrations. In contrast, much of the carbon dioxide in Earth's early atmosphere became part of the crustal materials, and the buildup of oxygen in Earth's atmosphere is a result of photosynthesis by plants. The development of Earth's habitable atmosphere, as contrasted with the torrid climate of Venus, appears to be directly related to Earth's distance from the Sun. Current analysis suggests that Earth's atmosphere would have evolved to the form found on Venus if the planet had been only 5 percent closer during the evolution of the atmosphere.

 On the remainder of the planets, the atmospheres appear to have retained the primordial nature associated with their formation. The air on Jupiter and Saturn, for example, is made up of nearly 100 percent diatomic hydrogen (H2) and helium (He), with small contributions of methane (CH4) and other chemical compounds. Much less is known regarding the atmospheres of the somewhat smaller Jovian planets Uranus and Neptune, although both are thought to be similar to those of Jupiter and Saturn.

      On both Jupiter and Saturn, colourful cloud bands and other regional phenomena that are located at different altitudes and latitudes circulate at speeds up to several hundreds of metres per second relative to each other. The large velocity shears associated with this motion create turbulent eddies on these planets—most notably Jupiter's Great Red Spot. The bright zones on these planets correspond to the tops of upwelling clouds in the cold upper atmosphere, whereas the more colourful bands correspond to the relatively warm lower atmosphere and may be associated with the occurrence of sulfur and phosphorus compounds. Both aurora displays and intense lightning have been observed on Jupiter and Saturn.

Roger A. Pielke, Sr.

Additional Reading

Introductory works
General references on meteorology, climatology, and aeronomy are provided in Ira W. Geer (ed.), Glossary of Weather and Climate: With Related Oceanic and Hydrologic Terms (1996). Introductory texts for meteorology and climatology include Edward Aguado and James E. Burt, Understanding Weather and Climate, 4th ed. (2007); Frederick K. Lutgens and Edward J. Tarbuck, The Atmosphere: An Introduction to Meteorology, 10th ed. (2007); C. Donald Ahrens, Meteorology Today: An Introduction to Weather, Climate, and the Environment, 8th ed. (2007); Roger G. Barry and Richard J. Chorley, Atmosphere, Weather, and Climate, 8th ed. (2003); Joseph M. Moran, Michael D. Morgan, and Patricia M. Pauley, Meteorology: The Atmosphere and the Science of Weather, 5th ed. (1997); Lee M. Grenci and Jon M. Nese, A World of Weather: Fundamentals of Meteorology, 4th ed. (2006); Richard A. Anthes, Meteorology, 7th ed. (1997); Eric W. Danielson, James Levin, and Elliot Abrams, Meteorology, 2nd ed. (2003); and Dennis L. Hartmann, Global Physical Climatology (1994). Books that portray the role of the atmosphere within the climate system include William R. Cotton and Roger A. Pielke, Sr., Human Impacts on Weather and Climate (2007); and P. Kabat et al. (eds.), Vegetation, Water, Humans, and the Climate: A New Perspective on an Interactive System (2004).

Thermodynamics, microclimates, and cloud processes
The behaviour of heat in the atmosphere is described in J.V. Iribarne and W.L. Godson, Atmospheric Thermodynamics, 2nd ed. (1981). The atmosphere near the Earth's surface (micrometeorology and microclimate) is discussed in depth by T.R. Oke, Boundary Layer Climates, 2nd ed. (1987, reprinted 1992); while atmospheric turbulence and the atmospheric boundary layer are presented in S. Pal Arya, Introduction to Micrometeorology, 2nd ed. (2001); Sheldon I. Green (ed.), Fluid Vortices (1995); J.R. Garratt, The Atmospheric Boundary Layer (1992); Zbigniew Sorbjan, Structure of the Atmospheric Boundary Layer (1989); and Roland B. Stull, An Introduction to Boundary Layer Meteorology (1988). Cloud processes (cloud microphysics), including precipitation formation, are given in William R. Cotton and Richard A. Anthes, Storm and Cloud Dynamics (1989); and William R. Cotton, Storms (1990).

Modeling atmospheric processes
The dynamics of the Earth's atmosphere are discussed mathematically in John A. Dutton, Dynamics of Atmospheric Motion (1995); José P. Peixoto and Abraham H. Oort, Physics of Climate (1992); and Murry L. Salby, Fundamentals of Atmospheric Physics (1996). The numerical modeling of atmospheric flow is described for larger-scale systems in T.N. Krishnamurti and L. Bounoua, An Introduction to Numerical Weather Prediction Techniques (1996); and for smaller-scale systems in Roger A. Pielke, Sr., Mesoscale Meteorological Modeling, 2nd ed. (2002).

Meteorological instrumentation
Weather instrumentation is described in Thomas P. De Felice, An Introduction to Meteorological Instrumentation and Measurement (1998). Remote sensing of the Earth's atmosphere by satellites is described in Graeme L. Stephens, Remote Sensing of the Lower Atmosphere: An Introduction (1994); and Stanley Q. Kidder and Thomas H. Vonder Haar, Satellite Meteorology: An Introduction (1995). The use of radars is presented in Richard J. Doviak and Dušan Zrnić, Doppler Radar and Weather Observations, 2nd ed. (1993, reissued 2006). The use of satellite and radar data in weather forecasting is described in M.J. Bader et al. (eds.), Images in Weather Forecasting: A Practical Guide for Interpreting Satellite and Radar Imagery (1995).Roger A. Pielke, Sr.

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Universalium. 2010.


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