Humidity

Humidity is the presence of water vapor in air, expressed through quantities that describe either its abundance or its proximity to saturation. Absolute humidity, specific humidity, and mixing ratio measure vapor content in different ways, whereas relative humidity compares actual vapor pressure with saturation vapor pressure at the same temperature. These distinctions explain why cooling air can raise its relative humidity without adding water vapor. (radiometrics.com)

Humidity connects the atmosphere with the water cycle and helps determine evaporation, cloud formation, and precipitation. Its distribution reflects exchanges with oceans and land, atmospheric circulation, and temperature. Water vapor also contributes to the greenhouse effect, making changes in atmospheric moisture important to climate and climate change. (ipcc.ch)

Humidity is measured using surface instruments, upper-air observations, and remote sensing. Beyond meteorology, it influences plant water use, human evaporative cooling, and the behavior of moisture-sensitive materials; its practical interpretation therefore depends on the quantity measured and the process being examined. (radiometrics.com)

Humidity Quantities and Units

Atmospheric humidity describes the gaseous water content of the atmosphere. Water vapor must be distinguished from liquid droplets and ice crystals suspended in clouds, and from water contained in soil or absorbed by materials. Humidity quantities describe either the amount of vapor present or its relation to saturation; they are not interchangeable measurements. In particular, a relative-humidity percentage does not specify the mass of vapor in a given volume of air. (forecast.weather.gov)

Absolute humidity is the vapor mass per volume of moist air, $\rho_v=m_v/V$, and is therefore a density. Its SI unit is $\mathrm{kg\,m^{-3}}$, although $\mathrm{g\,m^{-3}}$ is also used. Specific humidity, $q$, uses the combined mass of vapor and dry air as its denominator, whereas the water-vapor mixing ratio, $r$, uses only the dry-air mass:

Here $m_v$ and $m_d$ denote vapor and dry-air masses. Both quantities are dimensionless mass ratios, conventionally expressed in $\mathrm{kg\,kg^{-1}}$ or $\mathrm{g\,kg^{-1}}$; $10\ \mathrm{g\,kg^{-1}}$ equals $0.010\ \mathrm{kg\,kg^{-1}}$. Their different denominators mean that $q<r$ whenever vapor is present. (radiometrics.com)

Vapor partial pressure, $e$, is the contribution of water vapor to total atmospheric pressure, $p$. It is reported in pascals, hectopascals, or kilopascals, with $1\ \mathrm{hPa}=100\ \mathrm{Pa}$. Relative humidity, expressed as a percentage, is

where $e_s(T)$ is saturation vapor pressure at the air temperature $T$. The pressure units cancel, making relative humidity dimensionless: 50% means that actual vapor pressure is half the saturation value, not that half the air consists of water. The saturation reference must specify liquid water or ice, particularly below freezing. Precision calculations also account for the small pressure-dependent difference between saturation vapor pressure in moist air and that of pure water vapor. (weather.gov)

Under the ideal-gas approximation, vapor pressure and total pressure determine the mixing ratio:

The constant $\varepsilon\approx0.622$ is the ratio of the molar masses of water and dry air. Pressure quantities must use matching units, and $r$ must be inserted as a $\mathrm{kg\,kg^{-1}}$ ratio. For example, $e=2.00\ \mathrm{kPa}$ and $p=100.0\ \mathrm{kPa}$ give $r\approx12.69\ \mathrm{g\,kg^{-1}}$ and $q\approx12.53\ \mathrm{g\,kg^{-1}}$. Converting between $q$ and $r$ requires neither temperature nor pressure; converting either to $e$ requires total pressure. (weather.gov)

Absolute humidity additionally involves absolute temperature:

where $R_v$ is the specific gas constant for water vapor and $T$ is in kelvins. Thus converting vapor pressure to vapor density requires temperature, while converting relative humidity to vapor pressure requires temperature and a saturation-pressure relation. Obtaining specific humidity from relative humidity also requires total pressure. Expansion can change absolute humidity by changing volume even when vapor mass and dry-air mass remain unchanged; under those conditions, specific humidity and mixing ratio remain constant. (radiometrics.com)

Saturation, Temperature, and Phase Equilibrium

Saturation is a condition of thermodynamic equilibrium between water vapor and a condensed phase, not a literal limit on the amount of water that air can “hold.” Above a flat surface of pure liquid water at a fixed temperature, molecules continually escape through evaporation and return through condensation. At saturation, these opposing molecular fluxes balance, leaving no net transfer between liquid and vapor. The corresponding vapor partial pressure is the saturation vapor pressure, $e_s(T)$. Below it, liquid water tends to evaporate; above it, vapor tends to condense onto an available liquid surface. This equilibrium also exists without dry air: nitrogen and oxygen are not containers for water molecules. For precise measurements in moist air, small corrections account for interactions with other gases and the effects of total pressure. (prod-01-alb-www-noaa.woc.noaa.gov)

The temperature dependence follows the Clausius–Clapeyron relationship. Treating water vapor as an ideal gas and neglecting the liquid’s specific volume compared with that of the vapor gives

where $T$ is absolute temperature in kelvins, $L_v$ is the specific latent heat of vaporization, and $R_v$ is the specific gas constant for water vapor. The relationship implies an approximately exponential increase of saturation vapor pressure with temperature over modest temperature intervals. Near ordinary surface temperatures, the increase is roughly 7% per kelvin, although the fractional rate varies with temperature. The IPCC used this approximate scaling in its 2021 assessment of water-cycle changes. It describes a change in the equilibrium saturation pressure—not an automatic increase in actual vapor abundance. Maintaining unchanged relative humidity during warming requires additional water vapor; warming alone does not supply it. (ntrs.nasa.gov)

The condensed phase matters below freezing. Saturation vapor pressure over supercooled liquid water, $e_{sw}$, is higher than that over ice, $e_{si}$, at the same subfreezing temperature. Consequently, relative humidities referenced to liquid and ice differ:

For example, commonly used saturation-pressure formulations give approximately $e_{sw}=2.86\ \mathrm{hPa}$ and $e_{si}=2.60\ \mathrm{hPa}$ at $-10\,^\circ\mathrm{C}$, neglecting the small moist-air correction. Air at 100% relative humidity with respect to liquid water therefore has about 110% relative humidity with respect to ice. Conversely, air saturated over ice is only about 91% saturated over liquid water. A subfreezing humidity value must therefore identify its reference phase; standard meteorological relative humidity is commonly referenced to liquid water even below $0\,^\circ\mathrm{C}$. (eol.ucar.edu)

Supersaturation occurs when vapor pressure exceeds the saturation value for the specified phase. It is possible because equilibrium with an existing, flat surface is different from forming a new droplet or crystal. Creating a microscopic liquid droplet requires new surface area, and the curvature of a small droplet raises its equilibrium vapor pressure above that of a flat liquid surface. Dissolved substances have the opposite effect, lowering the equilibrium vapor pressure. Atmospheric particles such as sea salt and sulfate provide material on which droplets can develop; their activation as cloud condensation nuclei depends on particle size, composition, and ambient supersaturation. Thus, reaching 100% relative humidity over flat, pure water is not a universal threshold at which every particle becomes a cloud droplet. (aoml.noaa.gov)

Ice formation likewise has a nucleation barrier. Ice-nucleating particles can initiate freezing in supercooled droplets, with effectiveness depending on temperature and particle surface properties; mineral dust and some biological particles are documented examples. Saturation with respect to ice does not by itself guarantee the immediate appearance of ice crystals. Once crystals exist, vapor deposition can enlarge them whenever the surrounding air is supersaturated over ice, including conditions that remain subsaturated over liquid water. Supercooled droplets and ice crystals can therefore experience different directions of vapor exchange in the same air, reflecting their distinct equilibrium vapor pressures. (repository.library.noaa.gov)

Dew Point, Wet-Bulb Temperature, and Psychrometrics

The dew point, $T_d$, is the temperature at which air reaches saturation when cooled at constant pressure without changing its water-vapor mixing ratio. For ordinary atmospheric calculations, it satisfies $e_s(T_d)=e$, where $e$ is actual vapor pressure and $e_s$ is saturation vapor pressure over liquid water. Cooling toward the dew point removes sensible heat; it does not require adding water vapor. The frost point instead refers to saturation over ice, so the two reference temperatures differ below freezing. Dew-point depression, $T-T_d$, measures the difference between air temperature and dew point: at a given air temperature, a smaller depression indicates higher relative humidity. (sciencedirect.com)

The wet-bulb temperature describes cooling by evaporation, not cooling at unchanged vapor content. Water evaporating from a thermometer’s wetted wick consumes latent heat, lowering its temperature until heat supplied from the surroundings balances evaporative heat loss. The dry-bulb temperature is the ordinary air temperature measured without this evaporative cooling. In unsaturated air above freezing, dew point is normally below wet-bulb temperature, which is below dry-bulb temperature; at saturation, all three coincide. Thermodynamic wet-bulb temperature is defined by an ideal constant-pressure, adiabatic saturation process in which evaporation adds vapor to the air. A wetted thermometer approximates this quantity but involves simultaneous heat and mass transfer rather than the identical thermodynamic process. (fao.org)

Ventilation and atmospheric pressure must be considered when interpreting wet-bulb measurements. A commonly used psychrometric relation is $e=e_s(T_w)-Ap(T-T_w)$, with pressure $p$ and vapor pressures expressed in compatible units. FAO gives $A=0.000662\ \mathrm{K^{-1}}$ for an aspirated psychrometer with airflow around $5\ \mathrm{m\,s^{-1}}$, compared with $0.000800\ \mathrm{K^{-1}}$ for natural ventilation around $1\ \mathrm{m\,s^{-1}}$. Thus, a measured wet-bulb depression, $T-T_w$, cannot be converted into humidity independently of pressure and ventilation conditions. (fao.org)

Vapor-pressure deficit, $\mathrm{VPD}=e_s(T)-e$, expresses the departure of ambient air from saturation in pressure units, commonly kilopascals. It characterizes atmospheric evaporative demand but is not itself an evaporation rate. The relevant driving difference at a wet surface is approximately $e_s(T_s)-e$, where $T_s$ is surface temperature; ambient-air VPD equals this difference only when surface and air temperatures coincide. Available energy and resistance to vapor transport also govern the actual flux. Psychrometrics represents these relationships graphically: charts combine dry-bulb temperature, humidity ratio, relative-humidity curves, and wet-bulb and enthalpy lines. Each chart applies to a specified total pressure. ASHRAE’s 2017 SI charts, for example, include $101.325\ \mathrm{kPa}$ at sea level and $84.54\ \mathrm{kPa}$ for a nominal elevation of $1{,}500\ \mathrm{m}$. On a conventional temperature–humidity-ratio chart, moving at constant humidity ratio toward the saturation boundary identifies the dew point. (fao.org)

For a worked example, consider simultaneous observations of $T=25^\circ\mathrm{C}$ and relative humidity $50\%$. Using the liquid-water approximation employed in FAO’s Crop Evapotranspiration guidelines, published in 1998,

gives $e_s(25)\approx3.168\ \mathrm{kPa}$, hence $e\approx1.584\ \mathrm{kPa}$ and $\mathrm{VPD}\approx1.584\ \mathrm{kPa}$. Inverting the same expression gives

The dew-point depression is therefore approximately $11.1\ \mathrm{K}$. These calculated values assume saturation over liquid water and neglect small departures from ideal-gas behavior at ordinary atmospheric pressures. Subfreezing ice-reference calculations require an ice saturation relation; high-pressure applications require more complete moist-air property equations. Total pressure is additionally needed to locate the corresponding humidity ratio and wet-bulb temperature on a chart. (fao.org)

Sources, Sinks, and Atmospheric Transport

Atmospheric water vapor is supplied by evaporation from the ocean, lakes, rivers, wet soil, and intercepted water on vegetation, together with transpiration through plant leaves. Soil evaporation and plant transpiration are commonly grouped as evapotranspiration. These fluxes require energy for vaporization and a vapor-pressure difference between the evaporating surface and surrounding air. Radiation and heat transferred from the air provide energy; wind and turbulence carry vapor away, limiting its accumulation near the surface. Actual fluxes also depend on water availability and resistance to transport through soil and vegetation, so dry air alone does not guarantee rapid evaporation. (fao.org)

The coupling of energy supply and vapor transport underlies the combination method developed by Howard Penman in 1948 for estimating open-water evaporation. Evaporating one kilogram of water at approximately $20\,^\circ\mathrm{C}$ requires about $2.45\ \mathrm{MJ}$ of latent heat. Thus, evaporation represents both a moisture source and an energy transfer from the surface into the atmosphere. Over vegetated land, aerodynamic resistance governs transfer toward the air above the canopy, while surface resistance represents restrictions associated with stomata and the soil surface; agricultural applications are treated separately. (fao.org)

Condensation converts vapor into liquid droplets, whereas deposition converts it directly into ice. Both remove water from the gaseous phase, but neither necessarily removes water from the atmosphere: droplets and crystals can remain suspended in clouds. Precipitation reaching the surface transfers atmospheric water back to land or ocean. Falling rain that evaporates before reaching the ground instead returns water to the vapor reservoir. Likewise, evaporation of cloud droplets and sublimation of ice reverse earlier phase conversions. A vapor budget therefore distinguishes these internal exchanges from the net removal of atmospheric water across the surface boundary. (noaa.gov)

Horizontal advection transports vapor with the wind, while turbulent mixing redistributes it between neighboring air parcels and between the surface layer and air aloft. Convection carries moist lower-level air upward; larger-scale circulation organizes transport across regions and continents. Moisture convergence occurs where more water enters a region through atmospheric transport than leaves it, providing a supply additional to local evaporation. Atmospheric rivers are elongated, narrow corridors of concentrated vapor transport, whose strength depends on both moisture abundance and wind. During December 10–22, 2010, a series of these features fueled storms along the United States West Coast, with some locations receiving 11–25 inches, approximately 280–635 mm, of rain. (ipcc.ch)

Precipitable water measures the vapor mass above a unit area, rather than humidity at one height. Under hydrostatic balance,

where $q$ is specific humidity, $p_s$ is surface pressure, and $g$ is gravitational acceleration. Its units are $\mathrm{kg\,m^{-2}}$; $1\ \mathrm{kg\,m^{-2}}$ corresponds to approximately 1 mm of liquid water. It excludes cloud liquid and ice. Consequently, equal surface humidity need not imply equal column totals. (confluence.ecmwf.int)

Neglecting storage and transport of condensed water, the column moisture budget is

with $E$ positive upward evaporation, $P$ precipitation reaching the surface, and $\mathbf{v}_h$ horizontal wind. Negative flux divergence denotes convergence. If column storage changes little, precipitation can exceed local evaporation because transport supplies the difference. Conversely, moisture divergence can export locally evaporated water before it precipitates. An exact total-water budget also includes condensed-water storage and transport. (confluence.ecmwf.int)

Geographic, Vertical, and Temporal Patterns

Warm tropical air generally contains much more water vapor than cold polar air, even when the polar air has higher relative humidity. The difference reflects the strong temperature dependence of saturation vapor pressure, rather than a fixed water-holding capacity of air. As an illustrative calculation at an atmospheric pressure of $100\ \mathrm{kPa}$, air at $30^\circ\mathrm{C}$ and 80% relative humidity contains approximately $21\ \mathrm{g}$ of vapor per kilogram of moist air. Air at $-10^\circ\mathrm{C}$ and 90% relative humidity, referenced to ice, contains only about $1.5\ \mathrm{g/kg}$. These rounded values, calculated from standard humidity relationships, show why near-saturated polar air can have roughly one-fourteenth the vapor abundance of warm tropical air. (radiometrics.com)

Ocean–continent contrasts depend on both moisture supply and atmospheric circulation. The ocean provides an extensive evaporation source, whereas evaporation over land depends on soil moisture, vegetation, and available energy. Continental humidity therefore reflects imported oceanic moisture as well as local recycling through evapotranspiration. Distance inland alone does not determine humidity: prevailing winds can deliver moisture far into continents. Around $30^\circ$ north and south, descending branches of the large-scale circulation favor dry conditions. Descending air warms by compression, reducing relative humidity without necessarily removing additional vapor; limited rainfall can also restrict subsequent evaporation from land. (ipcc.ch)

Seasonal circulation changes can produce greater humidity contrasts than proximity to the coast. In a monsoon, changes in land–ocean heating and the pressure field reorganize winds and moisture transport. Summer inflow from the Indian Ocean brings humid conditions to much of South Asia, whereas winter continental outflow is generally drier. The North American monsoon typically develops over Mexico in June and extends into the southwestern United States in July, bringing moist air from the Pacific Ocean and Gulf of California. Mountains superimpose sharp local contrasts: rising air cools, condenses, and may lose water through precipitation on windward slopes; descending air warms on the lee side, lowering relative humidity and contributing to a rain shadow. Hawaiian northeast trade winds provide a clear example of wetter windward and drier leeward slopes. (nesdis.noaa.gov)

Vertically, most atmospheric vapor is concentrated in the lower troposphere, close to surface sources. Cooling with height and condensation during ascent generally reduce vapor abundance aloft, although individual profiles need not decline smoothly. A moist cloud layer can overlie a dry layer, and relative humidity can increase upward even while specific humidity decreases. Cold coastal waters can also sustain a cool, humid layer beneath a temperature inversion, as along California. Moisture entering the stratosphere is strongly constrained by the cold tropical tropopause: ice formation and sedimentation dehydrate ascending air. The tropical western Pacific is an especially important gateway, where low temperatures limit the vapor surviving passage into the stratosphere. (ipcc.ch)

Daily relative-humidity cycles often arise primarily from temperature changes. With vapor pressure approximately constant, relative humidity is usually highest near the daily temperature minimum and lowest during afternoon warmth. For example, a constant vapor pressure of $1.5\ \mathrm{kPa}$ gives about 88% relative humidity at $15^\circ\mathrm{C}$, but only 35% at $30^\circ\mathrm{C}$, using the FAO saturation-pressure approximation. No vapor removal is required for this decrease. Actual moisture content can nevertheless change through evaporation, transport, or dew formation; nighttime condensation may lower vapor pressure, while evaporation of dew after sunrise restores moisture. (fao.org)

Local humidity also depends on the surface and the depth of the air layer sampled. Vegetation and wet soil supply vapor, while shading changes temperature; their effects therefore involve both moisture addition and altered relative humidity. Observations across southern California coastal-to-desert environments have documented vegetation-related differences in atmospheric microclimates. Near cold water, cooling can raise relative humidity and produce a shallow humid layer without implying unusually large vapor abundance. California’s marine layer illustrates this distinction: cool, foggy coastal air may lie beneath warmer, clearer air only a short distance above or inland. (agris.fao.org)

Humidity in Weather Processes

Radiative cooling can bring surfaces and adjacent air to saturation without an increase in water-vapor content. On clear nights, the ground and exposed objects lose energy through infrared radiation. If a surface cools below the dew point of the air in contact with it, vapor condenses directly onto that surface as dew; the air at standard thermometer height need not be saturated. Fog, by contrast, consists of suspended droplets near the ground, rather than water deposited on objects. Radiation fog develops when cooling brings a sufficiently deep, moist near-surface layer to saturation, often under clear skies and light winds. California’s winter tule fog is an example: a cool, moist surface layer, clear skies above, and weak winds favor thick fog in the Central Valley. Clouds contain suspended droplets, ice crystals, or both, usually above the surface; dew, fog, and elevated cloud therefore describe different locations and forms of condensed water. (weather.gov)

Uplift produces saturation by a different mechanism: as an air parcel encounters lower pressure, it expands and cools. Air lifted along a front, over a mountain slope, or within a rising thermal can reach its condensation level and form cloud. Mixing can also produce saturation because saturation vapor pressure increases nonlinearly with temperature. Under an idealized constant-pressure approximation, equal mixtures of saturated air at $0\,^\circ\mathrm{C}$ and $20\,^\circ\mathrm{C}$ have a mean vapor pressure of about $14.8$ hPa, whereas saturation at the approximate mixture temperature of $10\,^\circ\mathrm{C}$ requires only $12.3$ hPa. Condensation removes the resulting excess vapor. Steam fog illustrates related processes: evaporation from warm water supplies moisture, and mixing with cold air produces smoke-like wisps above the water. Mixing with sufficiently dry air, however, can instead dissipate fog or cloud. (weather.gov)

Condensation releases latent heat into the rising air, partly offsetting its expansion cooling. The energy release is approximately $2.5\times10^6$ joules per kilogram of water condensed; this transfers energy previously absorbed during evaporation, rather than creating energy. Consequently, a saturated rising parcel generally cools more slowly than an unsaturated one. If it remains warmer and less dense than its surroundings, buoyancy sustains its ascent and can support the development of towering cumulus and thunderstorms. Moisture alone does not initiate this process: atmospheric instability and a lifting mechanism are also required, while a temperature inversion can inhibit ascent despite abundant low-level vapor. (training.weather.gov)

Dry-air entrainment—the incorporation of surrounding air into a cloud or updraft—promotes droplet evaporation, reducing cloud water and cooling the mixture. Cloud growth therefore depends partly on the balance between condensation and evaporation. Falling precipitation can likewise evaporate in an unsaturated layer beneath a cloud, moistening and cooling that layer. The cooling increases negative buoyancy and can strengthen downdrafts; strong descending air may spread outward at the surface as a downburst. If precipitation evaporates completely before reaching the ground, the visible falling streaks are called virga. Thus a radar echo aloft does not necessarily imply measurable surface rainfall. (weather.gov)

Forecasting uses the vertical distribution of moisture together with temperature, winds, and lifting mechanisms. On an atmospheric sounding, small separations between temperature and dew point identify nearly saturated layers, whereas large separations reveal dry layers that may inhibit cloud development or favor precipitation evaporation. Surface dew-point patterns indicate moist-air transport and moisture boundaries; persistent moisture convergence identifies regions where vapor accumulates and where thunderstorms may develop if instability and lift are also favorable. Surface relative humidity alone cannot establish rainfall probability: a shallow saturated layer may produce fog without rain, while precipitation generated in a moist layer aloft may fall through relatively dry surface air. The depth of moisture, stability, ascent, and evaporation along the precipitation’s path must therefore be assessed together. (weather.gov)

Instruments, Observation Networks, and Measurement Quality

Mechanical hygrometers translate moisture uptake into movement: a hair element lengthens as relative humidity rises, and a lever amplifies the change to move a pointer or recording pen. Horace-Bénédict de Saussure described a hair hygrometer in 1783. Wet- and dry-bulb psychrometers instead compare an ordinary thermometer with one covered by a wetted wick; evaporative cooling, ventilation, and atmospheric pressure determine the conversion from their temperature difference to humidity. Electronic instruments make automated recording possible: absorbed water changes a polymer dielectric’s capacitance, or a hygroscopic coating’s electrical resistance. Chilled-mirror instruments cool a reflective surface until an optically detected condensate layer can be maintained; the measured mirror temperature gives the dew or frost point, provided the condensate phase is identified. These methods measure different quantities directly, rather than interchangeable versions of the same signal. (collection-lhst.epfl.ch)

Surface observations require radiation shielding and adequate airflow so that sensor temperature represents the surrounding air. An inadequately ventilated enclosure can warm in sunlight and produce a low relative-humidity reading without any actual loss of atmospheric vapor. Near saturation, condensation on the sensing element can compromise its response and subsequent recovery; contamination and ageing require periodic checks. (weather.gov) Balloon-borne radiosondes extend observations vertically: the United States National Weather Service schedules routine soundings at 00:00 and 12:00 UTC. Aircraft supplement these profiles with measurements along flight routes; operational systems include capacitive sensors used by the MOZAIC and IAGOS programmes and the Water Vapour Sensing System II. Their sampling locations differ from those of fixed stations and balloon ascents. (weather.gov)

Response time is particularly important during ascent through sharp moisture gradients. A 2004 laboratory and field study developed time-lag corrections for Vaisala radiosonde humidity sensors using response measurements between $+25\,^\circ\mathrm{C}$ and $-60\,^\circ\mathrm{C}$; slower response at low temperatures can smooth or displace features in the observed profile. (journals.ametsoc.org) Solar heating introduces another error because a sensor warmer than ambient air registers lower relative humidity. Vaisala’s documented 2010 RS92 processing changes corrected both time lag and radiation heating, with the largest effects in daytime soundings approximately 10–15 km above the surface. Instrument model, processing version, and correction history therefore form part of the observation’s essential metadata. (vaisala.com)

Satellite remote sensing retrieves humidity from measured radiances rather than collecting an air sample. Infrared and microwave channels have different altitude sensitivities, allowing retrieval of vertical profiles, but the information is distributed across atmospheric layers rather than resolved at every reported output level. Opaque clouds prevent infrared sounders from profiling the atmosphere beneath them; companion microwave measurements provide information in cloudy conditions. EUMETSAT’s IASI products combine infrared observations with AMSU-A and MHS microwave measurements. (eumetsat.int) Ground-based global navigation satellite system observations, including GPS, exploit atmospheric signal delay. Surface pressure determines the hydrostatic contribution; subtracting it leaves a wet delay closely related to column-integrated water vapor. This provides a column measurement, not a surface relative humidity or a uniquely determined vertical moisture profile. (ngs.noaa.gov)

Laboratory calibration compares instruments with controlled humidity generators and reference hygrometers, establishing metrological traceability through documented standards and uncertainties. A generator may saturate gas at known temperature and pressure, then expand it or mix it with dry gas to establish a specified water mole fraction. Reference chilled-mirror instruments can transfer this calibration to field sensors, whose drift must be checked separately. (radiometrics.com) In their 2021 description of NIST’s Hybrid Humidity Generator, Christopher W. Meyer, Tobias Herman, and W. Wyatt Miller reported dew/frost-point generation from $-90\,^\circ\mathrm{C}$ to $+85\,^\circ\mathrm{C}$ at ambient pressure. The generator’s expanded measurement uncertainty, with coverage factor $k=2$, was below $0.025\,^\circ\mathrm{C}$ above $-60\,^\circ\mathrm{C}$. This reference uncertainty does not describe the complete uncertainty of an operational observation, which also depends on the calibrated instrument’s repeatability and measurement conditions. (nist.gov)

Climate Feedbacks, Trends, and Research Limitations

Water vapor absorbs terrestrial infrared radiation, contributing to the greenhouse effect. At otherwise unchanged temperatures, increased vapor reduces outgoing radiation and alters Earth’s energy budget. Warming increases saturation vapor pressure; if atmospheric moisture increases in response, its additional absorption amplifies the initial warming, producing a positive climate feedback. This differs from the externally imposed forcing associated with increasing concentrations of long-lived greenhouse gases, such as carbon dioxide: tropospheric water vapor responds strongly to temperature and is continually exchanged through evaporation, condensation, and precipitation. Upper-tropospheric moisture is particularly important to the radiative response, so surface humidity alone cannot characterize the strength of water-vapor feedback. (ipcc.ch)

The Intergovernmental Panel on Climate Change’s 2021 assessment reported a very likely increase in near-surface specific humidity over land and oceans since the 1970s, and a very likely decrease in relative humidity over much of the global land area since 2000, particularly at Northern Hemisphere mid-latitudes. Northern high latitudes showed increases rather than the widespread mid-latitude decline. The assessment expressed high confidence in the overall directions of increasing specific humidity and decreasing land relative humidity. Its principal surface-humidity figure covered 1973–2019. The slowdown in specific-humidity growth around 2000–2012 identified in the previous assessment had not persisted; the strong 2015–2016 El Niño contributed to unusually high surface moisture levels. These findings describe different quantities, not contradictory trends. (ipcc.ch)

Greater vapor abundance can coexist with declining relative humidity because relative humidity depends on vapor pressure relative to temperature-dependent saturation vapor pressure. If actual vapor pressure increases more slowly than saturation vapor pressure, relative humidity falls. Land generally warms more than the ocean, while moisture transported from the ocean supplies an important constraint on land humidity. Consequently, imported moisture can increase without matching the saturation-vapor-pressure increase over warmer land. Soil-water availability and vegetation responses, including changes in stomatal conductance, further modify local moisture supply. These mechanisms connect humidity changes to land–ocean warming contrasts rather than requiring an absolute decline in atmospheric water content. (ipcc.ch)

Future projections are distinct from these observations. In the 2021 assessment, CMIP6 climate models projected increases in global mean column-integrated water vapor of approximately 6–13% for 2041–2060 and 5–32% for 2081–2100, relative to 1995–2014, depending on the emissions scenario. The IPCC assigned high confidence to continued increases in global mean column water vapor and near-surface specific humidity over land, but medium confidence to region- and season-dependent decreases in land relative humidity. Regional moisture transport remains less constrained because circulation changes can reinforce or oppose thermodynamic changes. Cloud responses introduce additional uncertainty through spatially heterogeneous and microphysically complex processes. (ipcc.ch)

Long-term humidity records retain important uncertainties. Sparse sampling limits regional trend estimates, especially over Southern Hemisphere oceans; ship observation heights, instrument biases, and changes requiring record homogenization affect surface datasets. Reanalyses provide spatially complete estimates but remain sensitive to changes in their observing systems: apparent breaks in some column-water-vapor records coincide with such changes. These limitations complicate estimates of trend magnitude even where datasets agree on direction. Research questions also concern how convection and clouds redistribute upper-tropospheric moisture and how regional circulation changes modify humidity. Because upper-level moisture has disproportionate radiative importance, agreement in surface or column-total trends does not by itself establish that models reproduce its vertical distribution or cloud coupling. (ipcc.ch)

Plant Water Relations and Agricultural Uses

Transpiration is driven by the difference in water-vapor pressure between moist leaf interiors and surrounding air. Internal airspaces are commonly approximated as nearly saturated at leaf temperature, giving a leaf-to-air gradient of $e_s(T_{\mathrm{leaf}})-e_a$, where $e_s$ is saturation vapor pressure and $e_a$ is ambient vapor pressure. This differs from the vapor-pressure deficit calculated using air temperature: a sunlit leaf warmer than the air can experience a larger gradient. Vapor passes through stomatal pores and the boundary layer adjoining the leaf. Stomatal resistance and external transport resistance therefore restrict the flux; a large pressure difference does not, by itself, determine the rate of water loss. Near-saturation inside leaves is an approximation used in gas-exchange calculations, not an invariant condition. (academic.oup.com)

Radiation supplies energy for vaporization, while wind and turbulence transport vapor away from foliage. At canopy scale, aerodynamic resistance describes transfer from vegetation into the overlying atmosphere and depends on wind, measurement height, and canopy structure. Surface resistance represents the combined restrictions imposed by stomata, leaf area, and the evaporating soil surface. These resistances interact with water supply: when soil dries, increased stomatal resistance can reduce transpiration despite strong atmospheric evaporative demand. Radiation, temperature, and vapor-pressure deficit also influence stomatal behavior, with responses differing among crops and varieties. Thus, humidity measurements describe one component of the evaporative environment, rather than a direct measurement of actual plant water loss. (fao.org)

Agricultural calculations commonly use the FAO Penman–Monteith method to combine energy availability with aerodynamic vapor transport. Howard Penman’s 1948 combination approach joined energy-balance and mass-transfer calculations; a May 1990 FAO expert consultation recommended the Penman–Monteith approach as the standard for reference evapotranspiration. Its daily form is

Here $ET_0$ is in millimeters per day; $R_n$ and $G$ are net radiation and soil heat flux in megajoules per square meter per day; $T$ is mean air temperature in degrees Celsius; and $u_2$ is wind speed at 2 meters in meters per second. Vapor pressures are in kilopascals, while $\Delta$ and $\gamma$ are the saturation-vapor-pressure curve slope and psychrometric constant. The reference is hypothetical, adequately watered grass 0.12 meters high, with surface resistance $70\ \mathrm{s\,m^{-1}}$ and albedo 0.23. (fao.org)

Humidity observations enter chiefly through $e_a$, derived from dew point, wet- and dry-bulb measurements, or relative humidity paired with temperature. For daily calculations, FAO calculates $e_s$ from the average of saturation vapor pressures at daily maximum and minimum temperatures, rather than from mean temperature alone. Reference evapotranspiration is then converted to crop evapotranspiration under standard conditions using $ET_c=K_cET_0$, where the crop coefficient $K_c$ represents crop characteristics and growth stage. Water-stress calculations can separately reduce the transpiration component while retaining an explicit soil-evaporation component. Crop evapotranspiration is not identical to the irrigation requirement: effective rainfall, changes in stored soil water, and other water-balance contributions must also be considered. (fao.org)

Well-watered crops in hot, dry regions encounter both abundant energy and large vapor-pressure gradients. In humid tropical conditions, high ambient vapor pressure reduces the gradient even when radiation is substantial; increasing wind consequently tends to have a smaller effect than in dry air. Within a canopy, evaporation humidifies local air, and restricted air exchange can diminish the gradient around leaves. Such measurements represent a different microenvironment from standardized weather-station observations, normally taken at 2 meters over extensive grass or short vegetation. Crop-canopy humidity therefore cannot automatically replace reference-site data in evapotranspiration calculations without accounting for differences in exposure and measurement height. (fao.org)

Human Heat Exchange and Thermal Indices

Sweat contributes to human thermoregulation when it undergoes evaporation, drawing latent heat from the skin. Sweat production and evaporative heat loss are therefore distinct: liquid that remains on the skin or drips away does not provide the cooling associated with its evaporation there. The driving force for evaporation is the difference between water-vapor pressure at the wet skin surface and in the surrounding air. At otherwise comparable skin temperature, air movement, and clothing, increasing atmospheric vapor pressure reduces this difference and the potential rate of evaporative cooling. Relative humidity influences this process through its relationship to actual vapor pressure; a percentage by itself does not specify the skin-to-air gradient. (weather.gov)

Human heat exchange also includes convection, radiation, conduction, and respiratory losses. Air temperature determines the direction of convective exchange: air cooler than the skin can receive sensible heat from it, whereas hotter air can transfer heat toward the body. Solar radiation and infrared radiation from hot surroundings can add heat independently of humidity. Physical activity increases metabolic heat production, while clothing changes both insulation and resistance to vapor transfer. Impermeable protective garments, for example, can restrict sweat evaporation and create a warm, moist microenvironment close to the skin even when weather-station measurements indicate a comparatively drier atmosphere. Wind alters convective and evaporative transfer, so its effect cannot be represented by humidity alone. (cdc.gov)

The heat index expresses the combined influence of air temperature and relative humidity as an apparent temperature, rather than a thermometer measurement or a prediction of body temperature. Its physiological basis derives from Robert G. Steadman’s 1979 temperature–humidity analysis; the operational United States National Weather Service calculation uses a regression described by Lans P. Rothfusz in a 1990 technical attachment. A National Weather Service example gives a heat index of $124^\circ\mathrm{F}$, approximately $51^\circ\mathrm{C}$, for air at $100^\circ\mathrm{F}$, approximately $38^\circ\mathrm{C}$, and 55% relative humidity. At the same air temperature but 15% relative humidity, the reported index is $96^\circ\mathrm{F}$, approximately $36^\circ\mathrm{C}$. These values describe modeled apparent conditions; the measured air temperature remains $100^\circ\mathrm{F}$ in both cases. (wpc.ncep.noaa.gov)

Heat-index charts assume shade and light winds. The National Weather Service states that exposure to full sunshine can increase the apparent heat burden by up to $15^\circ\mathrm{F}$, equivalent to a temperature difference of about $8^\circ\mathrm{C}$; this is not a fixed correction for every sunny setting. The index does not explicitly incorporate actual solar exposure, local wind, workload, or the clothing worn by a particular person. Its calculation also has mathematical limits: the Rothfusz regression is not appropriate for apparent temperatures below about $80^\circ\mathrm{F}$, or $27^\circ\mathrm{C}$, where a simpler calculation is used, and it is not valid for extreme temperature–humidity combinations outside the underlying data range. (weather.gov)

Neither relative humidity nor dew point alone describes the complete thermal environment. Relative humidity depends on air temperature, while dew point identifies atmospheric vapor pressure without specifying air temperature, radiation, or air movement. Consequently, the same dew point can accompany different convective heat loads, and the same relative humidity can accompany different atmospheric vapor pressures. Heat-exchange assessments therefore distinguish environmental measurements from metabolic activity and clothing characteristics. A person performing strenuous work in protective coveralls and a lightly clothed person at rest can experience different heat loads under identical measured temperature and humidity. An individual humidity value is not a universal boundary between safe and unsafe exposure. (cdc.gov)

Buildings, Materials, and Humidity Control

Hygroscopic materials exchange water vapor with surrounding air until they approach an equilibrium moisture content determined by relative humidity, temperature, and their moisture-sorption history. This is water retained within the material, not the relative humidity of the adjacent air. Wood, paper, leather, and many textile fibers gain moisture as humidity rises and release it as humidity falls; their responses differ because of composition, coatings, thickness, and internal structure. The US Forest Service’s Wood Handbook, revised in 2021, describes how moisture changes produce dimensional movement in wood, with much greater movement across the grain than along it. In composite objects, unequal movement or restraint converts swelling and shrinkage into stress: wooden joints may loosen, paper may buckle, and paint or veneer layers may crack or separate from their supports. Diffusion and surface coatings delay equilibration, so a brief humidity fluctuation need not produce the same response as prolonged exposure. (research.fs.usda.gov)

Condensation occurs when a building surface is colder than the dew point of the air contacting it. Windows, cold-water pipes, and refrigerated air ducts can therefore become wet even when a room-average humidity measurement does not indicate saturation. The US Environmental Protection Agency’s moisture-control guidance gives an example of outdoor air with a dew point above 70 °F, approximately 21 °C, encountering an air-conditioning duct near 55 °F, approximately 13 °C: water condenses on the duct exterior. Before visible condensation occurs, cooling can already raise local relative humidity enough to favor deterioration. Conditions behind furniture or against exterior walls can consequently differ substantially from those at a centrally located sensor. Material composition and exposure duration also matter: an absorbent backing, an impermeable coating, and a bare metal surface respond differently to the same environment. (epa.gov)

Active humidity control changes the water-vapor balance. Refrigeration-based air conditioning and dehumidification cool air below its dew point, condense vapor, and remove the liquid; reheating can subsequently lower relative humidity without reversing that removal. Desiccants instead take up vapor by adsorption or absorption, allowing dehumidification without first cooling the air to saturation. Humidifiers add vapor through evaporation or steam delivery. Ventilation removes moisture when replacement outdoor air has a lower moisture content, but adds a humidity load when the incoming air has a higher dew point than indoor air. Heating alone lowers relative humidity without removing water vapor; cooling alone raises it until condensation begins. These distinctions explain why reaching a temperature set point does not necessarily mean that an air-conditioning system has controlled humidity adequately. (energy.gov)

In museums, enclosed display cases allow a collection-specific microclimate rather than requiring identical conditions throughout a building. Conditioned silica gel acts as a reversible moisture buffer, taking up water when humidity rises and releasing it when humidity falls. Its effectiveness depends on sorbent quantity, moisture capacity, enclosure leakage, and the difference between internal and external conditions. The Canadian Conservation Institute documents a case tested from March 1 to 11, 2021, with an air-exchange rate of 0.44 per day: without a sorbent, its internal relative humidity took about nine days to approach room conditions. Better sealing reduces the sorbent quantity and regeneration frequency needed, but does not itself dry damp contents. An enclosed object already containing excessive moisture can sustain a humid microclimate because its own moisture reservoir dries slowly. (canada.ca)

Requirements for cultural heritage differ according to the damage mechanism. Mold, involving fungi, depends on available moisture, substrate, temperature, and time; Canadian Conservation Institute experiments summarized in Technical Bulletin 26 report visible growth after roughly two weeks to three months at 75% relative humidity, compared with one to five days at 80–90%. Corrosion is strongly affected by surface contaminants: sodium chloride can deliquesce near 75% relative humidity, producing a corrosive liquid film. Mechanical damage, by contrast, depends particularly on the magnitude and duration of moisture-driven dimensional changes and on whether movement is restrained. Stefan Michalski’s conservation framework therefore distinguishes dampness, object-specific critical humidities, moisture-assisted chemical deterioration, and humidity fluctuations rather than treating one relative-humidity range as suitable for every collection. (canada.ca)

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