The atmosphere is a gravity-bound, mixed envelope of gases and suspended particles that envelops Earth. Its aerosols and particulates give rise to clouds and hazes and create the thin, blue limb visible from orbit above the tropospheric cloud tops. Viewed in clear space imagery, this atmospheric limb contrasts with celestial bodies such as the crescent Moon.
By dry-molar composition, air is dominated by nitrogen (78.08%) and oxygen (20.95%), with argon (0.93%), carbon dioxide (~0.04%) and trace gases; water vapor is highly variable, averaging roughly 1% at sea level and about 0.4% when averaged globally. The atmosphere’s total mass is approximately 5.15 × 10^18 kg, and because air density falls rapidly with height, roughly three quarters of that mass lies within the lowest ~11 km.
No sharp physical boundary separates air from outer space; nonetheless the Kármán line at 100 km is commonly adopted as the practical transition altitude. Vertical structure is therefore usefully described by layers defined by temperature gradients and physical properties: the troposphere (site of most weather and of air suitable for plant photosynthesis and animal respiration), the stratosphere (containing the ozone layer and a temperature inversion), the mesosphere, the thermosphere (associated with the ionosphere), and the outer exosphere. These vertical changes in composition, temperature and pressure control radiative transfer, chemical reaction rates and the ionization behavior of charged particles.
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The atmosphere performs multiple protective and regulatory functions: it causes most meteoroids to disintegrate before reaching the surface, attenuates ultraviolet radiation (notably via stratospheric ozone), reduces diurnal temperature extremes by mediating heat exchange, and retains heat through greenhouse processes that maintain surface conditions compatible with life.
Earth’s atmosphere has evolved substantially since its primordial accretion from the solar nebula. Geological and biological processes—including volcanic and tectonic outgassing, impacts, chemical weathering and, crucially, the emergence and activity of photoautotrophic organisms—have altered its composition and the global cycling of oxygen and carbon. In the modern era, human activities such as deforestation and combustion of fossil fuels have produced significant additional changes, driving climate warming and causing other perturbations like stratospheric ozone depletion and acid deposition.
The scientific study of the atmosphere (atmospheric science or aerology) spans climatology, atmospheric physics and related fields; paleoclimatology reconstructs past atmospheric states. Early instrumental pioneers in atmospheric layering and study include Léon Teisserenc de Bort and Richard Assmann.
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Composition
The standard composition of dry air is overwhelmingly dominated by three gases by molecular count: molecular nitrogen (N2) at 780,800 ppm (78.08% by volume; corresponding to ≈755,200 ppm or 75.52% by mass), molecular oxygen (O2) at 209,500 ppm (20.95% by volume; ≈231,400 ppm or 23.14% by mass), and argon (Ar) at 9,340 ppm (0.9340% by volume; ≈12,900 ppm or 1.29% by mass). These values define the conventional “dry‑air” reference abundances used in atmospheric work.
A set of minor and trace constituents occur at parts‑per‑million levels. Representative contemporary volume fractions include carbon dioxide (CO2) ≈412 ppm (≈626 ppm by mass), neon (Ne) 18.2 ppm (≈12.7 ppm by mass), helium (He) 5.24 ppm (≈0.724 ppm by mass), methane (CH4) 1.79 ppm (≈0.99 ppm by mass), and krypton (Kr) 1.14 ppm (≈3.3 ppm by mass). Trace species collectively constitute only a few hundred ppm by molecular count (≈0.0434% in the cited lower‑level representation) and include the primary greenhouse gases (CO2, CH4, N2O, O3), other noble gases (e.g., xenon) and numerous anthropogenic compounds. Measured concentrations of greenhouse gases such as CO2 and CH4 have been rising in recent decades, so snapshot ppm values should be interpreted as time‑dependent.
Water vapor is highly variable in both space and time and is therefore usually excluded from “dry‑air” composition tables. Averaged over the entire atmosphere, water vapor contributes roughly 0.25% of atmospheric mass, but its mole fraction near the surface ranges from order 10 ppm in very cold air to several percent (up to ≈5% in extremely hot, humid conditions). Tabulated non‑dry ranges are commonly reported as 0–30,000 ppm (0–3%) depending on local and temporal conditions.
Under typical atmospheric pressures the ideal gas law is accurate to within about 1%, so mole fraction and volume fraction are effectively equivalent; ppm therefore denotes parts per million by molecular count (mole or volume). Conversion between mole/volume fraction and mass fraction requires molecular weights: the mean molecular weight of dry air is ≈28.95 g·mol−1 (values cited ≈28.946–28.964 g·mol−1). This mean decreases with increasing humidity because water vapor (M ≈18.015 g·mol−1) is lighter than the average dry‑air molecule, a factor important for density and mass‑fraction calculations.
Vertically, turbulent mixing maintains nearly constant relative abundances of the well‑mixed gases through most of the atmosphere up to roughly 100 km. Above this altitude a transition zone (approximately 80–120 km) marks the decreasing importance of turbulence and the growing dominance of molecular diffusion; in the resulting heterosphere lighter species (e.g., He, H2) become relatively enriched with altitude. The MSIS‑E‑90 empirical model is commonly used to represent vertical compositional profiles, but it is applicable only above ≈85 km and therefore does not describe the lower, well‑mixed atmosphere.
Real ambient (unfiltered) air additionally contains variable amounts of aerosols and particulate matter of natural origin—mineral and organic dust, pollen, spores, sea spray, volcanic ash—and may contain industrial and anthropogenic gaseous and particulate pollutants (e.g., chlorine species, fluorinated compounds, elemental mercury vapor, hydrogen sulfide, SO2). Finally, users of compiled abundance tables should note analytical uncertainty: summing reported ppm values can exceed 1,000,000 ppm because of measurement and rounding errors (the referenced dataset sums ≈83.43 ppm above one million), indicating the limits of precision in compositional compilations.
A prism cross‑section is a conventional schematic for illustrating the vertical organization of Earth’s atmosphere: the major stratified layers are drawn to scale while local, transient features such as clouds or turbulence cells are not. Throughout the column both air pressure and density decline steadily with altitude, but temperature follows a more variable profile—decreasing, remaining nearly isothermal, or increasing in different altitude intervals—so the vertical temperature gradient (the lapse rate) serves as a primary diagnostic of atmospheric structure. The lapse rate, defined as the rate of change of temperature with height, exhibits a repeatable pattern that is routinely quantified by instrumented balloon soundings; observed lapse‑rate behavior provides a practical criterion for separating one atmospheric layer from another. On this basis the atmosphere is conventionally partitioned into five principal layers with typical altitude ranges: Troposphere (0–12 km), Stratosphere (12–50 km), Mesosphere (50–80 km), Thermosphere (80–700 km), and Exosphere (700–10,000 km).
Exosphere
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The exosphere is Earth’s outermost atmospheric region, so rarified that it is often considered a transition to interplanetary space rather than a conventional continuous atmosphere. Its lower boundary, the thermopause or exobase, marks the top of the thermosphere and rises with solar forcing from roughly 500 km to about 1,000 km. The outer boundary is not sharply defined: some definitions place it near 10,000 km, while others extend it to roughly 190,000 km—about halfway to the Moon—where Earth’s gravitational dominance weakens relative to other forces such as solar radiation pressure. The geocorona, a far-ultraviolet glow produced by neutral hydrogen, demonstrates the exosphere’s vast reach by extending at least 100,000 km from Earth.
Particle densities in the exosphere are extremely low and are dominated by atomic hydrogen with lesser amounts of helium and minute concentrations of carbon dioxide and newly formed oxygen nearer the exobase. Interparticle spacing is large enough that collisions are rare and the medium behaves as collisionless; individual atoms and molecules move on ballistic trajectories and may either fall back, migrate into or out of the magnetosphere, be picked up by the solar wind, or escape entirely to space. These escape processes produce measurable mass loss: roughly 3 kg of hydrogen and about 50 g of helium are lost per second, with much smaller fluxes of other species. Although meteorological phenomena do not occur at these altitudes, the exosphere contains many artificial satellites and thus forms the interface between engineered near‑Earth space and the broader interplanetary medium.
Thermosphere
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The thermosphere is the second‑highest major layer of Earth’s atmosphere, beginning at the mesopause near 80 km and extending upward to the thermopause (or exobase) roughly between 500 and 1,000 km, where it merges into the exosphere. Its upper boundary is highly variable: solar activity and the passage of the dawn–dusk terminator modulate the thermopause height and induce background density fluctuations of up to a factor of two, a dominant dynamical characteristic of this region.
Temperature, expressed as mean molecular kinetic energy, rises with altitude in the thermosphere and can reach values on the order of 1,500 °C at the top; this heating is driven mainly by absorption of solar extreme ultraviolet and X‑ray radiation. Despite these high kinetic energies, the gas is extremely rarified—typical molecular mean free paths are on the order of 1 km—so conductive heat transfer is negligible and the macroscopic sensation of “hot” does not apply. The thermosphere is free of clouds and water vapor, and classical hydrometeorological weather does not occur there, although electromagnetic and particle processes are active.
The thermosphere spatially overlaps the ionosphere (roughly 50–600 km), so ionization and plasma physics govern much of its behavior. Auroral displays commonly occur within this overlap (often near 100 km); auroral colors depend on the emitting species and altitude—green aurorae, for example, arise from excited atomic oxygen at about 120–400 km. Because many satellites, including the International Space Station (operating at ~370–460 km), orbit within the thermosphere, its low but variable neutral density together with enhanced plasma densities affects satellite drag, orbital decay rates, station‑keeping requirements, and radio propagation for systems interacting with ionized layers.
Mesosphere
The mesosphere is the third atmospheric layer, sandwiched between the stratosphere and the thermosphere, and extends from the stratopause at roughly 50 km to the mesopause near 80–85 km above mean sea level. Temperatures decline with height throughout this layer, reaching the mesopause—the coldest region of Earth—with mean values near 190 K (≈−85 °C).
Acoustic energy at audible frequencies is effectively absorbed before reaching mesosphere altitudes because atmospheric attenuation increases with the square of frequency; low-frequency infrasound can propagate upward but is difficult to generate at high power. Immediately below the mesopause, the very low water-vapor content can freeze into polar mesospheric (noctilucent) clouds— the highest clouds in the atmosphere—visible to the naked eye under twilight conditions when the Sun is approximately 4°–16° below the horizon. Lightning-driven upper-atmosphere discharges, grouped as transient luminous events, also occur episodically in this layer, producing brief optical emissions above tropospheric thunderstorms.
The mesosphere is the principal region for ablation of meteoroids and much low-orbit debris: incoming objects begin to decelerate and vaporize here, creating observable luminous trails. Atmospheric entry and reentry phenomena often reveal layered optical signatures—tropospheric afterglow appears orange, the stratosphere blue, and the mesosphere darker—and spacecraft commonly begin producing plasma or smoke trails upon encountering mesosphere densities. Because it lies above the operational ceiling of jets and conventional balloons yet below orbital altitudes, routine in situ access is limited; sounding rockets and specialized rocket-powered platforms are the primary means of direct investigation. Its intermediate altitude, extreme cold, distinctive cloud and electrical phenomena, and role in space–atmosphere interactions make the mesosphere a discrete and important domain for studies of atmospheric dynamics, chemistry, and optical processes.
The stratosphere is the second-lowest layer of Earth’s atmosphere, occupying the vertical interval between the tropopause (approximately 12 km above the surface) and the stratopause (roughly 50–55 km). It contains the ozone-rich band responsible for absorbing solar ultraviolet radiation; this absorption produces a temperature inversion—temperatures increase with altitude through much of the stratosphere—so that the layer is radiatively heated from above. Typical values illustrate this structure: temperatures near the tropopause reach about −80 °C (190 K), whereas the stratopause is substantially warmer and can approach 0 °C. Because roughly 99% of the atmosphere’s mass lies below 30 km and pressure at the top of the stratosphere is on the order of 1/1000 of sea-level pressure, air density falls sharply with altitude and dynamical activity is suppressed. The resulting strong thermal stratification yields a generally stable, low-turbulence environment largely free of the cloud systems and weather found in the troposphere; exceptionally cold conditions in the lower stratosphere can produce polar stratospheric (nacreous) clouds. The presence of a stratosphere on Earth owes to ozone chemistry and molecular oxygen; planets such as Mars and Venus lack comparable layers because of insufficient atmospheric oxygen and the consequent absence of an ozone-driven inversion. The stratosphere is nevertheless accessible to human activity: it is the highest layer routinely reached by jet aircraft, and extreme-altitude balloon operations have demonstrated human access to its lower reaches (notably Joseph Kittinger’s 1960 parachute jump from 31.3 km).
Troposphere
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The troposphere is Earth’s lowest atmospheric layer, averaging about 12 km in thickness and bounded above by the tropopause, whose altitude ranges from roughly 9 km at the poles to about 17 km at the equator and varies on short timescales with synoptic conditions. Heated primarily from the surface, the troposphere exhibits a characteristic lapse rate—temperature generally decreases with height—which keeps the lowest levels warmest and promotes vigorous vertical mixing (hence its name, from the Greek tropos, “turn”). This layer is comparatively dense because the weight of the overlying air compresses it, containing roughly 80% of the atmosphere’s mass and about 50% of that mass within the lowest 5.5 km. Nearly all atmospheric water vapor and associated moisture reside in the troposphere, with about 90% concentrated in the lower portion; consequently virtually all weather phenomena and cloud genera form within this layer. Strong convective systems occasionally penetrate the tropopause—very tall cumulonimbus clouds can intrude into the lower stratosphere—demonstrating episodic exchange between the layers. From space the troposphere produces a reddish filtering of sunlight and distinct cloud shadows, contrasting with the thin blue band of the stratosphere along the horizon. Human activity is largely confined to the troposphere: most conventional and propeller-driven aircraft operate within it, and persistent jet contrails form at ambient temperatures near −53 °C (≈−63 °F), typically encountered around 7.7 km altitude for modern engines.
The atmosphere is conventionally stratified into five principal temperature-defined layers—troposphere, stratosphere, mesosphere, thermosphere and exosphere—while additional secondary domains are recognized on the basis of composition, ionization, turbulent mixing and dynamical coupling. These secondary divisions overlap the main temperature layers and are important for understanding chemical processes, radiative effects and human uses of the near-space environment.
Ozone is concentrated in the lower stratosphere, roughly between 15 and 35 km, with a maximum near 32 km where local concentration reaches on the order of 15 parts per million. Although ozone remains a trace constituent relative to the major gases, about 90% of the planet’s atmospheric ozone resides in the stratosphere; the vertical extent and column abundance of this ozone-rich region vary with season and latitude.
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The ionosphere is the portion of the upper atmosphere ionized by solar radiation and is responsible for auroral displays, airglow and many space‑weather effects. By day it spans approximately 50–1,000 km, incorporating parts of the mesosphere, thermosphere and low exosphere; nocturnal recombination largely removes mesospheric ionization. The ionosphere forms the inner boundary of the plasmasphere and exerts a major control on long‑range radio‑wave propagation.
On the basis of turbulent mixing, the atmosphere is divided into the homosphere and heterosphere. The homosphere extends from the surface through the mesosphere into the low thermosphere and is chemically well mixed so that composition is essentially independent of molecular weight. This mixed regime terminates at the turbopause, near ~100 km, an altitude adopted by the Fédération Aéronautique Internationale as the conventional “edge of space,” about 20 km above the mesopause. Above the turbopause the heterosphere shows diffusive separation: heavier species (e.g., N2, O2) concentrate near its base while hydrogen dominates the uppermost region because molecular mean free paths exceed the scales of turbulent motions.
The planetary boundary layer is the surface‑coupled portion of the troposphere where turbulent diffusion links the atmosphere and Earth’s surface. It is typically well mixed during daytime but becomes stably stratified at night with weak or intermittent mixing; its depth is highly variable, from order 100 m on clear, calm nights to 1,000–1,500 m or more in convective afternoons.
The barosphere denotes the altitude range over which the barometric law (pressure decreasing predictably with height under Maxwellian velocity statistics) holds, extending from the surface up to the thermopause. Above the thermopause the particle velocity distribution departs from Maxwellian because high‑velocity atoms and molecules can escape, producing non‑Maxwellian behavior.
Global mean surface air temperature is commonly reported as approximately 14–15 °C (about 287–288 K), the precise value depending on the dataset and averaging methodology used.
Physical properties (standard reference and vertical structure)
The U.S. Standard Atmosphere (1962) furnishes a reference vertical profile of pressure, temperature and density referenced to fixed sea‑level base values: p0 = 101325 Pa, T0 = 288.15 K (15 °C) and ρ0 = 1.2250 kg·m−3. From these bases the model prescribes how pressure and density decline with altitude and how temperature varies through the layered atmosphere for engineering and geophysical comparisons.
Pressure and density fall rapidly with height: pressure decreases roughly exponentially from 101325 Pa at mean sea level to only a few pascals in the upper atmosphere, and density drops from 1.2250 kg·m−3 to orders of magnitude smaller by the upper mesosphere and thermosphere. These reductions govern aerodynamic forces, lift capability and aerodynamic heating for vehicles operating at different altitudes, and they make conventional airplane flight impossible once densities become extremely low.
The speed of sound a = sqrt(γRT) (with γ ≈ 1.4 and R ≈ 287 J·kg−1·K−1) is about 340.3 m·s−1 at standard sea level. Because a depends on the square root of absolute temperature, it generally decreases through the troposphere as temperature falls, then increases where temperature rises (notably in parts of the stratosphere and in the thermosphere), even while pressure and density continue to decline.
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Layered thermal structure in the standard profile. The troposphere extends from the surface to a tropopause near 11.0 km, characterized by a mean lapse rate of −6.5 K·km−1: T falls from 288.15 K at the surface to 216.65 K at 11.0 km, where p ≈ 22 632 Pa and ρ ≈ 0.364 kg·m−3. Above the tropopause the stratosphere (≈11–50 km) is approximately isothermal (~216.65 K) between ≈11–20 km in the standard model and then warms with height in the upper stratosphere, producing an increasing speed of sound despite continued density reduction. The mesosphere (≈50–85 km) again cools with altitude to the mesopause near ~85 km—the coldest region of the atmosphere—yielding very low densities that preclude aerodynamic lift. Above ~85 km the thermosphere exhibits extremely low particle densities but very large mean molecular kinetic temperatures; continuum‑flow assumptions used in aerodynamics break down here and orbital drag, rather than aerodynamic lift, controls vehicle dynamics.
Operational and geomorphological altitude markers. Human and operational reference altitudes plotted against the standard atmosphere include highest permanently inhabited settlements (~5 100 m), Mount Everest summit (8 848 m), typical commercial airliner cruise (≈10–12 km in the upper troposphere), high‑performance supersonic cruise (~18 km), high‑altitude reconnaissance/manned records (~25 km), and common stratospheric balloon operations (~20–40 km, with many soundings to 30–40 km). The Kármán line at 100 km is a conventional boundary between aerodynamic flight regimes and spaceflight; suborbital/sounding‑rocket apogees commonly lie in 50–200 km, the ISS orbits near 400 km where densities are extremely low but still generate measurable drag.
Orbital context and aerodynamic relevance. Satellite classifications relative to the standard atmosphere include LEO (~160–2 000 km), MEO (tens of thousands of kilometres for navigation satellites) and GEO at 35 786 km. At these altitudes aerodynamic lift is negligible; however, residual upper‑atmosphere density produces non‑zero drag in LEO that affects orbital decay. Thus, graphical comparisons of geometric altitude against density, pressure, temperature and speed of sound succinctly delineate regime boundaries: dense lower troposphere supports aerodynamic flight and control; aerodynamic performance diminishes through the stratosphere and mesosphere as density falls (even where temperature and sound speed rise); and above ∼100 km orbital mechanics, not aerodynamic lift, governs vehicle behavior.
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Geometric versus geopotential altitude. Standard‑atmosphere formulations and many engineering charts frequently employ geopotential altitude to compensate for the variation of gravity with height, while geometric altitude denotes true metres above mean sea level. For direct spatial correspondence to physical objects (aircraft, balloons, mountains, spacecraft) geometric altitude should be used when comparing their positions to pressure, density and temperature curves.
Pressure and thickness of Earth’s atmosphere
The International Standard Atmosphere sets mean sea‑level pressure at 101,325 Pa (commonly expressed as 1 atm, 760.00 Torr or 14.6959 psi), a reference used for measurements and calculations. Total atmospheric mass is about 5.1480×10^18 kg; this empirical value is roughly 2.5% smaller than the simple product of mean sea‑level pressure and Earth’s surface area because topographic highs remove air columns that would otherwise be counted in a flat‑Earth approximation.
Atmospheric pressure at any point is the weight of the air column above a unit area, so pressure varies in space and time with weather systems and local elevation. To first order pressure decreases exponentially with height: an increment of altitude equal to the scale height reduces pressure by a factor of e (≈2.718). On Earth the scale height is typically 5.5–6 km for the well‑mixed lower atmosphere (to roughly 80 km), although temperature, composition and local conditions alter this value. Above ~100 km the assumption of a single, well‑mixed gas breaks down and individual species follow their own scale heights; at mesospheric and thermospheric altitudes (≈200–300 km) the combined scale height rises to on the order of 20–30 km because of the changing temperature profile and molecular composition.
Mass is concentrated near the surface: about half of the atmosphere by mass lies below ≈5.6 km, roughly 90% below ≈16 km, and some 99.99997% below 100 km—the internationally recognized Kármán line commonly used to distinguish atmosphere from outer space and to classify high‑altitude travelers as astronauts. Representative human and vehicle altitudes illustrate this concentration: Mount Everest reaches 8,848 m; commercial airliners cruise near 9–12 km to exploit lower density and temperature for efficiency; stratospheric balloons ascend to ≈35 km; and the North American X‑15 attained 108.0 km in 1963.
Significant physical and engineering phenomena persist well above the Kármán line. Auroral activity and the initial incandescence of meteors occur above 100 km, and the ionospheric layers—critical for high‑frequency radio propagation—begin below 100 km and extend several hundred kilometres. The International Space Station operates roughly 370–460 km above Earth, a regime where residual atmospheric drag within the ionospheric F‑layer necessitates periodic reboosts; during heightened solar activity appreciable drag can affect satellites to altitudes of 600–800 km.
Because atmospheric density, composition and effective scale height change strongly with altitude and with solar and geomagnetic conditions, spacecraft and satellite operations must account for evolving aerodynamic drag, ionospheric radio effects and other space‑environment phenomena when selecting orbits, scheduling reboosts and planning communication strategies.
Temperature
In the lower atmosphere temperature generally declines with height: from sea level upward through the troposphere this negative vertical gradient (the lapse rate) persists until roughly the base of the stratosphere near 11 km. At that altitude the steady cooling with height gives way to a much weaker vertical gradient, and temperatures remain relatively constant over a substantial portion of the stratosphere. Above about 20 km—within the region commonly termed the ozone layer—the trend reverses and temperature increases with altitude because O2 and O3 absorb incoming solar ultraviolet radiation and convert it to heat, producing a positive temperature gradient. A second, more pronounced zone of increasing temperature occurs much higher, beginning in the thermosphere above ~90 km, where temperatures rise markedly with altitude. Separately, near-surface nocturnal conditions often produce a shallow temperature inversion (typically extending up to ~1,000 m): radiative cooling of the ground causes the surface to lose more energy than it receives, cooling the air in contact with the ground and causing temperature to increase with height over that layer.
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Speed of sound
In an ideal gas of fixed composition the local speed of sound depends only on the absolute temperature, T, and not independently on pressure or density; mathematically this is expressed as c = √(γRT), where γ is the ratio of specific heats and R the specific gas constant for air. Because c is controlled solely by T for a given composition, vertical variations in the speed of sound in the atmosphere follow the atmosphere’s temperature profile rather than the concomitant changes in pressure or density.
Near sea level under typical near-surface temperatures the speed of sound is approximately 340 m s⁻¹. In much colder layers such as the stratosphere (mean temperature ≈ −60 °C, ≈213 K) the speed of sound is substantially lower (on the order of 290 m s⁻¹), illustrating the sensitivity of c to temperature. Consequently, regions where temperature increases with height produce local increases in c, while regions of decreasing temperature with height produce decreases in c; therefore the vertical behaviour of acoustic propagation is governed primarily by the thermal stratification of the atmosphere.
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Density and mass
Atmospheric mass density declines rapidly with altitude and is characterized both locally (sea level) and globally (total column mass). Standard profiles such as the NRLMSISE-00 model supply temperature and mass-density as explicit functions of altitude; when plotted on logarithmic axes these profiles span many orders of magnitude, and visualization aids (auxiliary dotted lines at cubic values 8, 27, 64 … 729) are often used to make changes across decades easier to read.
At mean sea level air density is about 1.29 kg m−3 (equivalently 1.29 g L−1 or 0.00129 g cm−3). This value is not a direct measurement of mass but is calculated from observed pressure, temperature and humidity using the equation of state for air (a form of the ideal gas law), which relates those thermodynamic variables to mass density. Vertically, density is commonly approximated by the barometric formula—an exponential-like decrease governed by a scale height—sufficient for many engineering and meteorological applications. For high-precision needs (for example, modelling satellite drag and orbital decay in the thermosphere) more detailed, composition- and altitude-sensitive models such as NRLMSISE-00 are required.
The total mass of the global atmosphere is small relative to the solid Earth: commonly cited estimates are ≈5×10^15 tonnes (≈5×10^18 kg), or about 1/1,200,000 of Earth’s mass. A contemporary NCAR estimate gives a mean total atmospheric mass of 5.1480×10^18 kg, with interannual variability driven primarily by changes in atmospheric water vapor on the order of 1.2–1.5×10^15 kg depending on the data source (surface pressure versus direct water‑vapor measurements). NCAR’s component breakdown places mean water‑vapor mass at ≈1.27×10^16 kg and dry‑air mass at 5.1352 ± 0.0003 ×10^18 kg, separating the comparatively small but variable condensable fraction from the largely invariant non‑condensable background.
These numerical characterizations—sea‑level density, the barometric decrease with altitude, and total and constituent mass estimates—are foundational for practical and scientific calculations, from local buoyancy and aircraft performance to global mass‑balance assessments and precise predictions of low‑Earth‑orbit satellite decay.
Optical properties of the atmosphere govern how incoming solar shortwave radiation and outgoing terrestrial longwave radiation are partitioned within the climate system. Solar irradiance provides the dominant radiative input; the planet compensates by emitting energy at longer, infrared wavelengths. As radiation traverses the atmospheric column it is modified by radiative‑transfer processes—principally absorption, emission and scattering—so that both the amount and spectral character of energy reaching the surface or escaping to space are altered by atmospheric constituents.
Absorption and subsequent emission occur when gases, aerosol particles and hydrometeors (liquid droplets and ice crystals) take up radiation at specific wavelengths and then re‑radiate energy, typically at longer wavelengths; scattering redistributes photons directionally and can change the path length and apparent source of radiation without wholesale conversion to other spectral bands. The balance among these mechanisms determines whether radiation is transmitted to the surface, retained within the atmosphere, or redirected back to space.
Aerosols (for example, dust) and clouds are the principal reflectors in the troposphere. Cloud reflectivity depends sensitively on aerosol loading and cloud microphysics (drop and ice crystal size, phase and concentration); individual clouds can under some conditions reflect as much as ~70% of incident solar radiation. Globally, clouds are estimated to reflect about 20% of incoming solar energy and to account for roughly two‑thirds of Earth’s planetary albedo, making them the single largest control on planetary reflectivity and a key regulator of the global radiation budget.
Surface optical properties complete the radiative exchange: the fraction of solar energy not reflected or absorbed aloft reaches the surface and is partitioned between immediate reflection (surface albedo) and absorption followed by re‑emission as longwave infrared radiation. Thus surface characteristics feed back into vertical and global energy fluxes. Observational evidence of the importance of small‑scale microphysics is provided by satellite detections—e.g., twinkling glints from tropospheric ice crystals observed from space—demonstrating that microscale particle properties can produce measurable signals at very large observation distances.
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Scattering
When solar photons encounter Earth’s atmosphere their trajectories may be altered by interactions with molecules and particles; such redirection is termed scattering, while sunlight that arrives without such interactions is called direct radiation. Direct radiation produces a distinct solar image and well‑defined shadows because the beam retains its original directionality.
The portion of sunlight that has been scattered before reaching an observer is referred to as diffuse (or indirect) radiation. Under thick cloud cover or heavy aerosol loading most surface‑incident solar energy is diffuse, which homogenizes sky illumination and eliminates sharp shadows. A principal mechanism governing the sky’s colour is Rayleigh scattering, a wavelength‑dependent process whereby shorter wavelengths (blue) are scattered far more efficiently by atmospheric molecules and small particles than longer wavelengths (red). Consequently, regions of the sky away from the Sun appear blue.
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Solar elevation modulates the relative importance of these processes. When the Sun is near the horizon, sunlight traverses a much longer atmospheric path and undergoes proportionally greater removal of short wavelengths from the direct beam; the remaining transmitted light is therefore enriched in red and orange hues. Together, the balance between direct and diffuse radiation, wavelength‑dependent scattering, and solar zenith angle accounts for common diurnal optical phenomena: bright, sharply shadowed conditions under a strong direct beam at high solar angles; diffuse, shadowless illumination when scattering dominates (cloudy or aerosol‑laden skies); and pronounced reddening of sunrise and sunset light due to enhanced scattering of short wavelengths.
Absorption
Earth’s atmosphere does not transmit electromagnetic radiation uniformly; its transmittance varies strongly with wavelength because individual gases absorb specific spectral bands. The superposition of these molecular absorption features produces a composite atmospheric spectrum characterized by regions of high opacity interspersed with discrete low-opacity “windows” that permit radiation to pass relatively unimpeded between the surface and space.
Key molecular absorbers set the broad limits of these windows. Molecular oxygen and ozone together remove essentially all radiation at wavelengths shorter than ~300 nm, eliminating the ultraviolet-C portion of the solar spectrum from reaching the surface. Water vapor is the dominant absorber at many longer wavelengths, producing pronounced attenuation throughout much of the near- and mid-infrared, particularly above ~700 nm. The precise positions and strengths of absorption bands depend on gas concentrations, pressure, and temperature, so the effective transparency of any window is determined by the combined atmospheric state.
At the microscopic scale, absorption converts photon energy into internal energy of molecules, producing local heating; conversely, the atmosphere cools by re-emitting radiation at other wavelengths. In ground-based astronomical observations, atmospheric absorption of celestial signals—termed telluric contamination—must be corrected for to recover true source spectra.
The best-known transmission region, the optical window, opens near the ~300 nm ozone cutoff, contains the visible band (~400–700 nm), and extends into the near-infrared to roughly 1100 nm, allowing much of the Sun’s shortwave radiation to reach the surface and enabling optical and near-IR astronomy. Beyond this, partial infrared windows occur where absorption is weaker, and a separate radio window exists at much longer wavelengths. On Earth the radio window spans approximately from ~1 cm up to ~11 m in wavelength, defining the range of radio frequencies that penetrate the atmosphere with relatively low attenuation and thereby supporting ground-based radio astronomy and telecommunications.
Emission
Emission denotes the radiative loss of energy from a body and complements absorption; its spectral shape and total power follow a black‑body distribution such that warmer bodies emit more energy and at shorter wavelengths (Wien’s law and the Stefan–Boltzmann relation). The Sun, with an effective temperature near 6000 K, therefore radiates most strongly around 500 nm in the visible band, whereas the much cooler Earth surface (roughly 290 K on average) emits peak radiation near 10 µm in the longwave infrared, a region invisible to the human eye.
Because the atmosphere itself has temperature, it also emits infrared radiation, and this atmospheric emission is a key component of the surface energy balance. Emissive behavior depends on gas composition and vertical temperature structure: water vapor and cloud liquid are especially effective absorbers and emitters in the thermal infrared. Consequently, clear nights permit stronger net radiative cooling of the surface than cloudy nights, since clouds intercept and re‑emit longwave radiation back toward the surface. The same principle explains enhanced nocturnal cooling at high elevations, where thinner, cooler air and reduced downward infrared flux from overlying atmosphere and clouds allow more rapid loss of terrestrial longwave energy.
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The greenhouse effect arises because certain atmospheric constituents (notably H2O and CO2) are largely transparent to incoming solar shortwave radiation but absorb and re‑emit outgoing terrestrial longwave radiation, sending a portion of that energy back toward the surface. This radiative insulation raises Earth’s mean surface temperature from an estimated −18 °C in the absence of greenhouse gases to the observed global average of about 15 °C, illustrating the central climatic role of longwave atmospheric absorption and emission.
Refractive index
Air has a refractive index only marginally above unity, so light propagates slightly slower through the atmosphere than through a vacuum and its direction is susceptible to even small spatial variations in that index. Systematic gradients—most importantly persistent vertical changes—produce continuous bending of rays; although each incremental deviation is small, the integrated effect along long optical paths can appreciably shift apparent directions and positions of distant objects.
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Because refractive bending can occur in the same sense as the Earth’s curvature, atmospheric stratification can cause rays to follow a curved trajectory that extends an observer’s optical line-of-sight beyond the geometric horizon, enabling objects to be seen that lie just below the true horizon. Temperature is the dominant environmental control on air’s refractive index in the lower atmosphere, so regions with strong vertical temperature gradients generate the largest refraction effects; modest temperature differences near the surface therefore suffice to produce substantial optical bending.
An extreme manifestation of temperature-gradient–driven refraction is the mirage: smoothly varying layers of air continuously refract light and produce displaced, stretched, inverted or multiple images of distant features and surfaces. Similar stratified refraction also alters the apparent shapes of celestial bodies at low elevation; for example, solar outlines near the horizon appear flattened and otherwise distorted by passage through the layered atmosphere.
Circulation
Atmospheric circulation in the troposphere is the principal mechanism by which heat is redistributed across the planet, working in concert with ocean currents. Its large-scale architecture is set mainly by the latitudinal contrast in incoming solar radiation and by Earth’s rotation, which together determine the preferred directions and scales of flow; while interannual variability alters details, the fundamental pattern is relatively stable. Seasonal changes in the latitude of maximum solar heating, driven by Earth’s axial tilt, shift the circulation belts through the year, and the uneven placement of continents and oceans perturbs and fragments the idealized zonally symmetric circulation, producing pronounced regional deviations.
A convenient conceptual model of the meridional overturning comprises three latitude-banded cells. Near the equator the Hadley cell is powered by intense convective ascent of warmed air; that rising air diverges aloft and moves poleward before descending and returning equatorward near the surface. The mid-latitude Ferrel cell is a thermally indirect, transitional circulation between the Hadley and Polar cells: its mean surface flow tends poleward while the upper-level return flow is equatorward, and the cell’s maintenance depends strongly on dynamical exchanges with transient eddies and storms. At high latitudes the Polar cell completes the meridional overturn, with vertical exchanges and poleward upper-level flow closing the loop between mid- and polar latitudes.
Sharp vertical and horizontal shear at the interfaces between these cells concentrates winds into narrow, fast-moving jets that generally run west to east at roughly 9,100 m altitude. These jets shift position seasonally—intensifying in winter when equator-to-pole temperature gradients are greatest—and their instabilities organize and steer the synoptic-scale weather systems that dominate the mid-latitudes. Superimposed on this mean overturning are wave and tidal disturbances generated by uneven solar heating and the diurnal cycle: small-scale gravity waves propagate momentum into the upper atmosphere, planetary-scale Rossby waves govern large-scale meanders of the flow, and atmospheric tides act as periodic oscillations that transport energy upward from the troposphere and stratosphere into higher layers.
Earliest atmosphere (Hadean)
During the Hadean eon Earth’s primordial envelope was inherited from the surrounding solar nebula and was therefore dominated by molecular hydrogen with abundant simple hydrides (e.g., H2O vapor, CH4, NH3) rather than the oxygen- and nitrogen-rich mixture of the modern atmosphere. In composition and character this first atmosphere more closely resembled the hydrogen-rich envelopes of present-day gas giants than a terrestrial, oxidized air.
Sustained high-energy bombardment in the Hadean, including frequent large meteorite impacts, injected sufficient heat into the near-surface environment and atmospheric column to drive substantial volatile loss through thermal escape and atmospheric stripping. The Moon-forming giant impact—commonly attributed to a collision with a Mars-sized body (Theia)—was especially consequential, melting and ejecting large portions of Earth’s mantle and crust, altering planetary mass distribution and contributing to lunar accretion.
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The energetic disruption of mantle and crustal material promoted extensive outgassing: molten and fractured rock released large quantities of water vapor and other volatiles into the atmosphere. As the post-impact planet cooled, this steam condensed and precipitated, accumulating as the first sizable bodies of liquid water and giving rise to the early oceans by the end of the Hadean.
In sum, early Earth experienced a transition from a nebula-derived, hydrogen–hydride atmosphere to an impact-modified, steam-dominated regime that lost many of its most volatile constituents, enabled ocean formation, and produced the coupled Earth–Moon system through large-scale mantle and crustal ejection.
Second atmosphere
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The transition from the Hadean to the Archean marked a fundamental reorganization of Earth’s surface and atmosphere. Progressive solidification of the crust curtailed advective heat transfer, allowing atmospheric water vapor to condense and precipitate a global ocean; this shift reorganized surface hydrology and the modes of heat exchange between lithosphere, ocean and atmosphere. Continued volcanic outgassing, together with volatiles delivered and redistributed by large impacts during the Late Heavy Bombardment, generated an early Archean atmosphere dominated by N2 with substantial CO2, CH4 and noble gases, providing the chemical background for subsequent ocean–atmosphere interactions.
Rapid exchange between atmosphere, hydrosphere and crust established major geochemical cycles early. Dissolution of CO2 into surface waters and chemical weathering of crustal rocks mobilized Ca and Mg, producing carbonate minerals that accumulated as sediments; carbonate-bearing deposits are preserved as far back as ~3.8 Ga. Carbon isotope relationships in ancient materials indicate that key features of the modern carbon cycle — production, burial and weathering pathways — were already operating by ~4.0 Ga.
By roughly 3.4 Ga the atmosphere had evolved into a comparatively stable, nitrogen‑dominated “second atmosphere.” This compositional shift coincides with some of the earliest biological signatures: paleobiological evidence points to phototrophic and other microbial communities emerging by ~3.5 Ga, implying that biology began to modulate atmospheric composition and redox-sensitive geochemical cycles soon after the second atmosphere was established. Over the Neoarchean a long interval of cyanobacterial photosynthesis gradually elevated free oxygen in surface environments; stromatolite assemblages dated to ~2.7 Ga document active phototrophic communities, and this prolonged biological oxygen production ultimately contributed to the Great Oxygenation Event in the late Archean–early Proterozoic.
Climate records of the Archean present the well‑known “faint young Sun” paradox: despite a Sun ≈30% dimmer than today, geological evidence indicates broadly clement surface conditions with persistent liquid water, punctuated by at least one major glacial episode centered near ~2.4 Ga. Sedimentary sequences also preserve high temporal variability in redox conditions: detailed paleoredox records from Gabon (≈2.15–2.08 Ga) record dynamic oxygenation changes likely driven by large perturbations in the global carbon and oxygen cycles (for example the Lomagundi‑Jatuli carbon isotope excursion), highlighting the complex interplay among biology, carbon burial and atmospheric composition during this interval.
Third atmosphere
Plate tectonics has governed the long‑term chemical evolution of Earth’s atmosphere by cycling carbon between the air and large continental carbonate reservoirs; this tectonic redistribution of CO2 modulates global carbon budgets and thereby constrains the redox trajectory of the atmosphere over geological time. A pivotal transformation occurred in the early Proterozoic: around 2.4 billion years ago oxygen produced by cyanobacterial photosynthesis first accumulated persistently in the atmosphere. The geological signature of this Great Oxygenation Event is the disappearance of banded iron formations, reflecting exhaustion of the abundant surface reductants (ferrous iron, reduced sulfur species and methane) that had previously scavenged nascent O2.
Atmospheric oxygenation proceeded only after net O2 production exceeded the capacity of these chemical sinks; thereafter O2 levels fluctuated markedly through the Proterozoic, including prolonged intervals of ocean euxinia (widespread anoxic, sulfide‑rich conditions), before rising toward a more stable oxidizing state by the close of the Precambrian. The diversification and expansion of eukaryotic photoautotrophs—notably green and red algae—provided additional pulses of oxygen, particularly following Cryogenian glaciations, and set the stage for the Ediacaran emergence of complex multicellular animals. During the Phanerozoic the oxygen record continued to vary: the Cambrian diversification of more active metazoans coincided with increasing O2, and the Carboniferous reached a pronounced maximum (near ~35% O2) compared with the modern ~21%.
Two primary biological controls drive long‑term O2 change: increases in global carbon fixation associated with the spread and innovation of terrestrial and aquatic plants, and countervailing biological consumption by animals and by plant respiration and photorespiration. Geological processes also modulate atmospheric O2—oxidative weathering and decomposition of sulfide minerals consume oxygen, while volcanic degassing injects CO2 that can indirectly stimulate photosynthetic O2 production even as volcanism supplies sulfur species that remove O2. Empirically, episodes of elevated oxygen commonly coincide with accelerated metazoan innovation, implicating O2 availability as a key environmental precondition for ecological and physiological complexity. Nonetheless, the precise balance of biological and geological drivers and the detailed timing of O2 excursions remain incompletely resolved.
Air pollution
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Air pollution denotes the release of gases, aerosols, particulate matter and biological agents into the atmosphere at concentrations that harm human health, ecosystems or material assets. Rapid population growth together with industrialization and widespread motor vehicle use have greatly expanded emissions, producing visible and adverse outcomes such as urban smogs, acid deposition and pollution-related respiratory and cardiovascular diseases. Certain anthropogenic compounds, notably chlorofluorocarbons and other ozone-depleting substances, have degraded the stratospheric ozone layer that normally attenuates biologically harmful ultraviolet radiation.
Since about 1750, and accelerating after the Industrial Revolution, human activities have elevated atmospheric concentrations of the principal greenhouse gases—carbon dioxide (CO2), methane (CH4) and nitrous oxide (N2O)—and altered land cover through deforestation and wetland drainage. The combined effect of increased greenhouse forcing and land‑use change has produced measurable global warming: average surface temperatures in 2011–2020 were about 1.1 °C higher than around 1850. In the Northern Hemisphere troposphere, CO2 displays a strong seasonal cycle driven by the terrestrial biosphere: concentrations typically peak in spring (around May) and decline after the summer peak in vegetation productivity as photosynthetic uptake draws down atmospheric CO2. This seasonal drawdown is especially pronounced over high‑latitude boreal forests where the summer uptake of carbon is large.
The rise in greenhouse gases and persistent air pollutants has a range of geographically significant consequences. Thermal expansion of warming oceans and melting of land ice are driving sea‑level rise; greater CO2 uptake is acidifying surface waters; glaciers and snowpacks are retreating, jeopardizing water supplies for downstream communities; and warmer, drier conditions contribute to more frequent and severe extreme weather events and wildfires. Together these changes increase the risk of widespread ecological disruption and mass mortality in both terrestrial and marine systems.