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Atmospheric physics
Atmospheric physics
Atmospheric physics
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Atmospheric sciences
Meteorology
Climatology
Aeronomy
Glossaries

Within the atmospheric sciences, atmospheric physics is the application of physics to the study of the atmosphere. Atmospheric physicists attempt to model Earth's atmosphere and the atmospheres of the other planets using fluid flow equations, radiation budget, and energy transfer processes in the atmosphere (as well as how these tie into boundary systems such as the oceans). In order to model weather systems, atmospheric physicists employ elements of scattering theory, wave propagation models, cloud physics, statistical mechanics and spatial statistics which are highly mathematical and related to physics. It has close links to meteorology and climatology and also covers the design and construction of instruments for studying the atmosphere and the interpretation of the data they provide, including remote sensing instruments. At the dawn of the space age and the introduction of sounding rockets, aeronomy became a subdiscipline concerning the upper layers of the atmosphere, where dissociation and ionization are important.

Remote sensing

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Brightness can indicate reflectivity as in this 1960 weather radar image (of Hurricane Abby). The radar's frequency, pulse form, and antenna largely determine what it can observe.
Main article: Remote sensing

Remote sensing is the small or large-scale acquisition of information of an object or phenomenon, by the use of either recording or real-time sensing device(s) that is not in physical or intimate contact with the object (such as by way of aircraft, spacecraft, satellite, buoy, or ship). In practice, remote sensing is the stand-off collection through the use of a variety of devices for gathering information on a given object or area which gives more information than sensors at individual sites might convey.[1] Thus, Earth observation or weather satellite collection platforms, ocean and atmospheric observing weather buoy platforms, monitoring of a pregnancy via ultrasound, magnetic resonance imaging (MRI), positron-emission tomography (PET), and space probes are all examples of remote sensing. In modern usage, the term generally refers to the use of imaging sensor technologies including but not limited to the use of instruments aboard aircraft and spacecraft, and is distinct from other imaging-related fields such as medical imaging.

There are two kinds of remote sensing. Passive sensors detect natural radiation that is emitted or reflected by the object or surrounding area being observed. Reflected sunlight is the most common source of radiation measured by passive sensors. Examples of passive remote sensors include film photography, infrared, charge-coupled devices, and radiometers. Active collection, on the other hand, emits energy in order to scan objects and areas whereupon a sensor then detects and measures the radiation that is reflected or backscattered from the target. radar, lidar, and SODAR are examples of active remote sensing techniques used in atmospheric physics where the time delay between emission and return is measured, establishing the location, height, speed and direction of an object.[2]

Remote sensing makes it possible to collect data on dangerous or inaccessible areas. Remote sensing applications include monitoring deforestation in areas such as the Amazon Basin, the effects of climate change on glaciers and Arctic and Antarctic regions, and depth sounding of coastal and ocean depths. Military collection during the Cold War made use of stand-off collection of data about dangerous border areas. Remote sensing also replaces costly and slow data collection on the ground, ensuring in the process that areas or objects are not disturbed.

Orbital platforms collect and transmit data from different parts of the electromagnetic spectrum, which in conjunction with larger scale aerial or ground-based sensing and analysis, provides researchers with enough information to monitor trends such as El Niño and other natural long and short term phenomena. Other uses include different areas of the earth sciences such as natural resource management, agricultural fields such as land usage and conservation, and national security and overhead, ground-based and stand-off collection on border areas.[3]

Radiation

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This is a diagram of the seasons. In addition to the density of incident light, the dissipation of light in the atmosphere is greater when it falls at a shallow angle.
See also: Radiation and Effect of sun angle on climate

Atmospheric physicists typically divide radiation into solar radiation (emitted by the sun) and terrestrial radiation (emitted by Earth's surface and atmosphere).

Solar radiation contains variety of wavelengths. Visible light has wavelengths between 0.4 and 0.7 micrometers.[4] Shorter wavelengths are known as the ultraviolet (UV) part of the spectrum, while longer wavelengths are grouped into the infrared portion of the spectrum.[5] Ozone is most effective in absorbing radiation around 0.25 micrometers,[6] where UV-c rays lie in the spectrum. This increases the temperature of the nearby stratosphere. Snow reflects 88% of UV rays,[6] while sand reflects 12%, and water reflects only 4% of incoming UV radiation.[6] The more glancing the angle is between the atmosphere and the sun's rays, the more likely that energy will be reflected or absorbed by the atmosphere.[7]

Terrestrial radiation is emitted at much longer wavelengths than solar radiation. This is because Earth is much colder than the sun. Radiation is emitted by Earth across a range of wavelengths, as formalized in Planck's law. The wavelength of maximum energy is around 10 micrometers.

Cloud physics

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Main article: Cloud physics

Cloud physics is the study of the physical processes that lead to the formation, growth and precipitation of clouds. Clouds are composed of microscopic droplets of water (warm clouds), tiny crystals of ice, or both (mixed phase clouds). Under suitable conditions, the droplets combine to form precipitation, where they may fall to the earth.[8] The precise mechanics of how a cloud forms and grows is not completely understood, but scientists have developed theories explaining the structure of clouds by studying the microphysics of individual droplets. Advances in radar and satellite technology have also allowed the precise study of clouds on a large scale.

Atmospheric electricity

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Cloud-to-ground lightning in the global atmospheric electrical circuit
Main article: Atmospheric electricity

Atmospheric electricity is the term given to the electrostatics and electrodynamics of the atmosphere (or, more broadly, the atmosphere of any planet). The Earth's surface, the ionosphere, and the atmosphere is known as the global atmospheric electrical circuit.[9] Lightning discharges 30,000 amperes, at up to 100 million volts, and emits light, radio waves, X-rays and even gamma rays.[10] Plasma temperatures in lightning can approach 28,000 kelvins and electron densities may exceed 1024/m3.[11]

Atmospheric tide

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Main article: Atmospheric tide

The largest-amplitude atmospheric tides are mostly generated in the troposphere and stratosphere when the atmosphere is periodically heated as water vapour and ozone absorb solar radiation during the day. The tides generated are then able to propagate away from these source regions and ascend into the mesosphere and thermosphere. Atmospheric tides can be measured as regular fluctuations in wind, temperature, density and pressure. Although atmospheric tides share much in common with ocean tides they have two key distinguishing features:

i) Atmospheric tides are primarily excited by the Sun's heating of the atmosphere whereas ocean tides are primarily excited by the Moon's gravitational field. This means that most atmospheric tides have periods of oscillation related to the 24-hour length of the solar day whereas ocean tides have longer periods of oscillation related to the lunar day (time between successive lunar transits) of about 24 hours 51 minutes.[12]

ii) Atmospheric tides propagate in an atmosphere where density varies significantly with height. A consequence of this is that their amplitudes naturally increase exponentially as the tide ascends into progressively more rarefied regions of the atmosphere (for an explanation of this phenomenon, see below). In contrast, the density of the oceans varies only slightly with depth and so there the tides do not necessarily vary in amplitude with depth.

Note that although solar heating is responsible for the largest-amplitude atmospheric tides, the gravitational fields of the Sun and Moon also raise tides in the atmosphere, with the lunar gravitational atmospheric tidal effect being significantly greater than its solar counterpart.[13]

At ground level, atmospheric tides can be detected as regular but small oscillations in surface pressure with periods of 24 and 12 hours. Daily pressure maxima occur at 10 a.m. and 10 p.m. local time, while minima occur at 4 a.m. and 4 p.m. local time. The absolute maximum occurs at 10 a.m. while the absolute minimum occurs at 4 p.m.[14] However, at greater heights the amplitudes of the tides can become very large. In the mesosphere (heights of ~ 50 – 100 km) atmospheric tides can reach amplitudes of more than 50 m/s and are often the most significant part of the motion of the atmosphere.

Aeronomy

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Main article: Aeronomy
Representation of upper-atmospheric lightning and electrical-discharge phenomena

Aeronomy is the science of the upper region of the atmosphere, where dissociation and ionization are important. The term aeronomy was introduced by Sydney Chapman in 1960.[15] Today, the term also includes the science of the corresponding regions of the atmospheres of other planets. Research in aeronomy requires access to balloons, satellites, and sounding rockets which provide valuable data about this region of the atmosphere. Atmospheric tides play an important role in interacting with both the lower and upper atmosphere. Amongst the phenomena studied are upper-atmospheric lightning discharges, such as luminous events called red sprites, sprite halos, blue jets, and elves.

Centers of research

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In the UK, atmospheric studies are underpinned by the Met Office, the Natural Environment Research Council and the Science and Technology Facilities Council. Divisions of the U.S. National Oceanic and Atmospheric Administration (NOAA) oversee research projects and weather modeling involving atmospheric physics. The US National Astronomy and Ionosphere Center also carries out studies of the high atmosphere. In Belgium, the Belgian Institute for Space Aeronomy studies the atmosphere and outer space. In France, there are several public or private entities researching the atmosphere, as an example météo-France (Météo-France), several laboratories in the national scientific research center (such as the laboratories in the IPSL group).

See also

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References

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Further reading

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Revisions and contributorsEdit on WikipediaRead on Wikipedia

Atmospheric physics

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Atmospheric physics is a branch of that applies fundamental principles of physics to understand and model the behavior, structure, and processes of Earth's atmosphere, as well as those of other planetary atmospheres. It focuses on phenomena occurring across a wide range of spatial and temporal scales, from molecular interactions to global circulation patterns, using tools such as equations, budgets, and energy transfer models to explain atmospheric motions, energy balances, and constituent interactions. Central to atmospheric physics are key subfields including atmospheric thermodynamics, which governs transformations and transfers in dry and moist air parcels, influencing formation and stability; atmospheric , which examines how solar and terrestrial is absorbed, scattered, and emitted by atmospheric gases, aerosols, and clouds, playing a critical role in the and regulation; and atmospheric dynamics, which studies fluid motions, wave propagations, and using principles from and hydrodynamics to predict phenomena like storms, jet streams, and planetary boundary layers. Additional notable aspects include and physics, which investigate microphysical processes such as droplet formation, growth, and particle scattering that affect precipitation, visibility, and ; and , encompassing charge separation in thunderstorms, propagation, and the global electric circuit, with implications for production and ionospheric interactions. These areas are interconnected, requiring integrated observational techniques—from satellite remote sensing to in-situ measurements—and computational modeling to address challenges like parameterizing small-scale processes for large-scale and predictions.

Atmospheric Structure and Composition

Vertical Layers

The Earth's atmosphere is vertically stratified into distinct layers primarily based on variations in with altitude, a that influences global patterns, dynamics, and interactions. These layers are defined by thermal boundaries rather than sharp discontinuities, with transitions occurring over several kilometers. The primary divisions include the , , , , and , each exhibiting unique temperature lapse rates and physical characteristics that govern atmospheric stability and energy distribution. The , extending from the surface to approximately 12 km altitude (varying from 8 km at the poles to 18 km in the ), is the lowest layer where phenomena predominantly occur. in this layer decreases with altitude at an average environmental of about 6.5 K/km, driven by adiabatic cooling of rising air parcels. decreases exponentially with height, following a of roughly 8 km, while air density drops from about 1.2 kg/m³ at to around 0.2 kg/m³ at the . This layer contains the bulk of the atmosphere's mass, accounting for over 80% of its total, and its boundary, the , acts as a stable cap inhibiting deep . Above the lies the , spanning roughly 12 to 50 km altitude, where temperature increases with height due to absorption of ultraviolet radiation—a process briefly involving stratospheric . This temperature inversion creates a stable layer with a that can reach positive values up to 2-3 K/km near the stratopause. continues to decline exponentially, reaching about 0.008% of sea-level values (approximately 80 Pa) at 50 km, and falls to approximately 0.001 kg/m³, supporting phenomena like polar stratospheric clouds. The stratopause marks the upper boundary at around 50 km, separating it from the overlying . The , from about 50 to 85 km, is the coldest atmospheric layer, with temperatures plummeting to as low as -90°C at the due to minimal solar heating and by . The here is steeply negative, averaging -2 to -3 K/km, while pressure drops to around 0.001% of surface levels and density to about 10^{-5} kg/m³, conditions that lead to the formation of noctilucent clouds during summer at high latitudes. This layer serves as a transitional zone where atmospheric drag affects satellites and predominantly burn up. Extending from 85 to approximately 600 km is the , where temperatures rise dramatically with altitude, often exceeding 1,000 due to absorption of high-energy solar radiation, though the low means actual heat content is minimal. The is positive and variable, with pressure nearing levels (approximately 2 \times 10^{-5} Pa at 200 km) and as low as 10^{-12} kg/m³ at the upper edge, enabling auroral displays and ionospheric interactions. The thermopause at around 600 km delineates its boundary. Beyond 600 km, the represents the tenuous outermost layer, transitioning gradually into interplanetary , with no well-defined upper boundary. Here, particle densities are so low (around 10^{-15} kg/m³ or less) that collisions are rare, and temperatures are not meaningfully defined due to the sparse medium; atoms and molecules can escape into , contributing to atmospheric loss over geological timescales. This layered structure is fundamentally maintained by , expressed by the equation dPdz=−ρg,\frac{dP}{dz} = -\rho g, where PP is atmospheric pressure, zz is altitude, ρ\rho is air density, and gg is gravitational acceleration (approximately 9.8 m/s² near the surface). This balance between the weight of the overlying air column and pressure gradients results in the observed exponential decrease in pressure and density with height. The identification of these vertical layers dates to the early 20th century, when balloon-borne and rocket soundings by scientists like Léon Teisserenc de Bort and Robert Goddard revealed temperature inversions and gradients through direct measurements of pressure and temperature profiles. Modern understanding has been refined by satellite missions, such as NASA's TIMED (Thermosphere Ionosphere Mesosphere Energetics and Dynamics) spacecraft, launched in 2001, which provided high-resolution data on upper atmospheric structure and dynamics.

Chemical Constituents

The chemical composition of Earth's atmosphere is primarily gaseous, with dry air consisting of approximately 78.08% nitrogen (N₂), 20.95% oxygen (O₂), 0.93% argon (Ar), and 0.04% carbon dioxide (CO₂) by volume, alongside trace gases such as neon (Ne), helium (He), and methane (CH₄) that collectively make up less than 0.01%. Water vapor, which is highly variable and can range from nearly 0% in arid regions to up to 4% in humid tropical areas, is not included in the dry air composition but significantly influences atmospheric properties like density and heat capacity. Vertically, the composition remains relatively uniform below about 100 km altitude due to turbulent mixing in the , where molecular nitrogen and oxygen dominate as the primary constituents. Above this level, in the heterosphere, begins to separate gases by , leading to an increase in lighter species and the dissociation of O₂ into atomic oxygen (O), whose concentration rises sharply in the due to by . Isotopic variations, such as the ratio of ¹⁸O to ¹⁶O in atmospheric oxygen and , provide key proxies for paleoclimate reconstruction, as processes during and reflect past and hydrological cycles. Human activities have notably altered concentrations; for instance, atmospheric CO₂ has risen from pre-industrial levels of about 280 parts per million (ppm) to approximately 427 ppm as of November 2025, primarily due to combustion and . Similarly, (CH₄) levels have increased by over 150% since pre-industrial times, with current levels at approximately 1920 ppb as of November 2025, —particularly and paddies—accounting for roughly 40% of current anthropogenic emissions. These constituents determine key physical properties, including the mean molecular weight of dry air, which is approximately 28.97 g/mol, influencing in convective processes and the rates of gases. The governs atmospheric behavior as PV=nRTPV = nRT, where PP is , VV is , nn is the number of moles, RR is the , and TT is temperature; incorporating molar mass MM yields the density equation ρ=PMRT\rho = \frac{PM}{RT}, which highlights how composition affects air and thus vertical stability.

Core Physical Processes

Thermodynamics

In atmospheric physics, of thermodynamics, stated as dU=δQ−δWdU = \delta Q - \delta W, where UU is , δQ\delta Q is added to the system, and δW\delta W is work done by the system, ensures during atmospheric processes such as compression and expansion of air parcels. In adiabatic processes, where no exchange occurs (δQ=0\delta Q = 0), this law implies that rising air cools due to expansion work, maintaining energy balance without external sources. The second law of thermodynamics introduces , which increases in irreversible processes like the mixing of air masses with different temperatures or compositions, quantifying the atmosphere's tendency toward disorder and limiting reversible heat engines in natural cycles. Heat transfer mechanisms in the atmosphere are dominated by , particularly in the , where updrafts driven by transport vertically over large scales, far outweighing other modes in influencing patterns. Conduction plays a minor role due to air's low thermal conductivity of approximately 0.026 W/m·K at 20°C, limiting direct flow between adjacent air molecules except near the surface. A critical process is the release of during , providing about 2.5×1062.5 \times 10^6 J/kg, which warms rising moist air parcels and enhances convective instability. To analyze atmospheric stability, potential temperature θ\theta is employed, defined by the formula θ=T(P0P)R/Cp,\theta = T \left( \frac{P_0}{P} \right)^{R / C_p}, where TT is the actual , PP is , P0P_0 is a reference pressure (typically 1000 hPa), RR is the specific gas constant for dry air (287 J/kg·K), and CpC_p is the at constant pressure (1004 J/kg·K); this quantity remains constant during dry adiabatic processes, allowing comparison of air parcels at different pressures. An atmosphere is if θ\theta increases with height, neutral if constant, and unstable if decreasing, providing a thermodynamic basis for predicting vertical motion tendencies. The dry adiabatic , derived from the first law for unsaturated air rising without heat exchange, is Γd=g/Cp≈9.8\Gamma_d = g / C_p \approx 9.8 K/km, where gg is (9.8 m/s²), representing the rate of decrease with altitude in a neutrally dry atmosphere. The Clausius-Clapeyron equation describes the relationship between temperature and saturation vapor pressure, given by dPvdT=LTΔV,\frac{dP_v}{dT} = \frac{L}{T \Delta V}, where PvP_v is the saturation vapor pressure, LL is the latent heat of vaporization, TT is temperature, and ΔV\Delta V is the change in specific volume between vapor and liquid phases; this relation limits atmospheric humidity by showing how warmer air can hold exponentially more water vapor. Historically, Émile Clapeyron derived an early form in 1834 from thermodynamic principles, which Rudolf Clausius refined in 1850 to its modern integrated version, establishing a foundational tool for understanding phase changes in meteorological contexts. This equation underpins the role of thermodynamics in cloud formation by constraining moisture availability for condensation.

Fluid Dynamics

Fluid dynamics in atmospheric physics governs the motion of air masses, treating the atmosphere as a compressible, rotating fluid subject to forces like pressure gradients, , and . The fundamental equations describing these motions are the Navier-Stokes equations, which express the conservation of momentum for a viscous fluid. In simplified form for atmospheric applications, they are written as ρ(DvDt)=−∇P+ρg+μ∇2v,\rho \left( \frac{D \mathbf{v}}{Dt} \right) = -\nabla P + \rho \mathbf{g} + \mu \nabla^2 \mathbf{v}, where ρ\rho is , v\mathbf{v} is , PP is pressure, g\mathbf{g} is , and μ\mu is dynamic , with the DDt\frac{D}{Dt} accounting for advective changes. These equations are often further approximated using the Boussinesq assumption, which neglects density variations except in the term, treating ρ\rho as constant to simplify calculations for flows where density changes are small compared to the , such as in convective motions. This approximation is valid when the relative density perturbation Δρ/ρ≪1\Delta \rho / \rho \ll 1 and the flow depth is much less than the of the atmosphere. A key feature of atmospheric is the influence of , introduced via the , which acts on moving air parcels perpendicular to their velocity. The Coriolis parameter is f=2Ωsin⁡ϕf = 2 \Omega \sin \phi, where Ω\Omega is Earth's (7.29×10−57.29 \times 10^{-5} s−1^{-1}) and ϕ\phi is , varying from zero at the to a maximum at the poles. In large-scale flows, this leads to geostrophic balance, where the balances the : 1ρ∇P=fk×v\frac{1}{\rho} \nabla P = f \mathbf{k} \times \mathbf{v}, resulting in winds parallel to isobars with speed vg=1ρf∣∇P∣v_g = \frac{1}{\rho f} |\nabla P|. This balance approximates mid-latitude synoptic-scale circulations, where rotational effects dominate over . Turbulence in the atmosphere arises due to high Reynolds numbers, defined as Re=ρVLμ\mathrm{Re} = \frac{\rho V L}{\mu}, where VV is a and LL is a scale; flows become when Re≳103\mathrm{Re} \gtrsim 10^3, which is typical for atmospheric motions given air's low and scales from meters to thousands of kilometers. Near the surface, modifies the geostrophic flow within the , modeled by Ekman layer , where the layer depth is approximately δE=2KMf\delta_E = \sqrt{\frac{2 K_M}{f}}
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