• 1.1 Atmospheric structure
• 1.2 Gas laws and basic atmospheric forces
• 1.3 Atmospheric pressure gradient and buoyancy
• 1.4 Atmospheric tides
• 1.5 Atmospheric moisture
• 1.6 Evaporation and latent heat

1.1 Atmospheric structure

1.1.1 Temperature-related layers

There are four temperature-related atmospheric regions. The outermost is the thermosphere, within which the temperature rises rapidly with height until about 300 km above Earth's surface. In parts of the thermosphere, the temperature varies diurnally (daily) by 30% or so (200 °C – 300 °C ), due to absorption of ultra-violet solar radiation as thermal energy, without the ability to re-radiate. Depending on the sunspot activity cycle, theoretical temperatures at the 150–300 km level vary between 200 °C and 1700 °C but due to the rarified atmosphere there is little sensitive heat capacity. The absorbed heat is conducted downward below 100 km where the atmosphere can re-radiate at night.

Temperature decreases rapidly with height in the mesosphere(from the Greek 'mesos' — middle); the minimum of about –90 °C is reached at the mesopause located at about 80 km where atmospheric pressure is about 0.01 hPa. Carbon dioxide in the mesosphere is an important absorber of terrestrial infra-red radiation. A group of wind systems is centred within the mesosphere, just above the stratopause, extending into the stratosphere and, to some extent, the thermosphere.

Most of the atmosphere's ozone [O₃] is contained within the stratosphere (from the Latin 'stratum' — layered); the O₃ is produced between the 30 and 60 km levels by reaction between atomic oxygen [O] and molecular oxygen [O₂]. Atmospheric circulation transports ozone down to the 25 km level where maximum density occurs — this is the ozone layer. The ozone content tends to concentrate at lower levels in the higher latitudes during the winter months and is transported to lower latitudes during spring. Ozone blocks about 90% of the sun's UV radiation — roughly all radiation between 0.25 and 0.35 micrometres. That UV energy absorption results in the temperature in the upper half of the stratosphere increasing until the stratopause. The temperature in the lower half of the stratosphere tends to remain constant or increase slightly with height, thus the layer is usually very stable. Some vertical mixing occurs and there is east-west and west-east circulation, but once gases or particles enter the stratosphere they tend to stay in it for long periods.

The troposphere (from the Greek 'tropos' — [over]turning) — its thickness varying from about 8 km at the poles to 28 km at the equator, and varying daily and seasonally — contains virtually all the atmospheric water and more than 90% of the air mass. Condensation of water vapour, forming clouds, occurs almost exclusively in the lowest 8 km where the water vapour comprises up to 3% or 4% of the atmosphere by volume. The troposphere is heated by terrestrial long-wave radiation plus turbulent mixing of latent and sensible heat. Vertical air movement can be pronounced and temperature decreases linearly with height until the tropopause. The low temperature at the tropopause (–40 °C to –50 °C in the mid-latitudes) allows very little water vapour to pass above it; refer atmospheric moisture below.

1.1.2 Composition-related layers

The troposphere, stratosphere and mesosphere constitute the homosphere (from the Greek 'homos' — same) in which the composition of the atmosphere is more or less uniform throughout. The composition is primarily nitrogen (78%), oxygen (21%) and argon (<1%), plus other trace gases and particles; the two major non-permanent gases O₃ and H₂O, plus CO₂, are particularly important as radiation absorbers because of their triatomic structure. The average atmospheric relative molecular mass throughout the homosphere is about 29 kg per 1000 moles. (A mole is the basic SI unit of amount of substance. One mole of any substance contains 6 x 10²³ molecules, the latter being the number of molecules in 12 grams of carbon-12.)

The composition changes above the mesopause. The atmospheric gases tend to separate into layers according to the relative molecular weight of the individual components, thus the average relative molecular mass decreases with height. This second composition layer, which extends to inner space, is the heterosphere (from the Greek 'heteros' — other). The relative molecular weight of the main atmospheric components is:

There is little or no nitrogen above 200 km, atomic oxygen dominates between 300 and 1000 km, helium between 1000 and 2000 km, and hydrogen above that.

1.1.3 Radiation-related layers

In the photochemical ionosphere (which is mostly contained within the thermosphere but also partly extends into the neighbouring mesosphere), cosmic radiation of high-energy sub-atomic particles and the absorption of much of the solar ultraviolet radiation separates negative electrons from oxygen and nitrogen molecules. The ions and free electrons move rapidly under the influence of electrical forces — the ionospheric wind — and the ionosphere is highly conductive; see the global circuit. Oxygen is chemically active when affected by shortwave ultraviolet radiation and molecular (diatomic) oxygen, O₂ , dissociates into atomic (monatomic) oxygen. Above 150 km the molecular nitrogen separates out owing to its higher mass, and the atmosphere is predominantly atomic oxygen. The excitation of oxygen and nitrogen atoms by collision with charged particles (separated hydrogen electrons and protons) from outburst emissions of solar wind produces the aurorae in the ionosphere.

Several ionisation layers are formed in the ionosphere that affect radio communications:

• The F2 or Appleton layer is at about 400 km by day, descending to 200 km at night. Ionisation varies from 10⁶ free electrons/cc during the day to 10⁵ at night. The layer refracts LF, MF and HF waves. But VHF, UHF and higher frequencies are not significantly affected.
  
  
• The F1 layer is at about 200 km. Nitrogen is ionised by short-length UV radiation in the F layers.
  
  
• The E or Heaviside-Kennelly layer is at 90–150 km. It has 10⁵ free electrons/cc by day, but disappears at night. The E layer partially reflects LF, MF, HF and sometimes VHF signals back to earth. At night, it is replaced by the F2 layer at 200 km. The longest X-rays ionise oxygen and nitrogen.
  
  
• The D layer, where N₂O is ionised by medium length UV, exists only during daylight at 50–90 km. It reflects LF and VLF waves, absorbs MF and attenuates HF. Solar outbursts (sunspots, flares) radiate X-rays in abnormal quantity and ionise the E and D layers strongly, lowering their altitude and adversely affecting HF communications during the day.
  
  
• The changes in the ionisation layers affect the sky waves of navigation aids such as non-directional beacons. Errors in directional indications will increase, particularly during the morning and evening twilight periods.

The energy-absorbing region from the tropopause to the D layer, i.e. the stratosphere and the mesosphere, is the ozonosphere. Ultraviolet radiation dissociates the water vapour that reaches the stratosphere and higher regions into hydrogen and oxygen atoms.

When such atoms reach the exosphere (from the Greek 'exo' — outside) — above the thermopause at about 500 km and extending out for an indeterminate distance, where the lighter components predominate — some atoms, particularly helium and hydrogen, will reach escaping velocity. The temperature within the exosphere remains roughly constant with height, although it varies daily and seasonally.

The magnetosphere limits the Earth's geomagnetic field. Within it are the Van Allen belts of high-energy solar wind and cosmic radiation particles trapped by the magnetic field. The outer, mainly electron, belt is centred about 18 000 km above the equator. The inner, more energetic and mainly proton, belt is centred at 3000 km. Changes within the magnetosphere may influence weather.

1.2 Gas laws and basic atmospheric forces

The density (the mass of a unit of volume) of dry air is about 1.225 kg/m³ at mean sea level [msl] and decreases with altitude. The random molecular activity within a parcel of air exerts a force in all directions and is measured in terms of pressure energy per unit volume, or static pressure. This activity, i.e. the internal kinetic energy, is proportional to the absolute temperature. (Absolute temperature is expressed in kelvin units [K]. One K equals one degree Celsius and zero degrees in the Celsius scale is equivalent to 273 K.) There are several gas laws and equations that relate temperature, pressure, density and volume of a gas.

Boyle's law:

At a constant temperature the volume [V] of a given mass of gas is inversely proportional to the pressure [P] upon the gas; i.e. PV = constant.

The pressure law:

At a constant volume the pressure is directly proportional to temperature [T] in Kelvin units.

Charles' law:

At a constant pressure gases expand by about 1/273 of their volume, at 273 K, for each one K rise in temperature; i.e. the volume of a given mass of gas at constant pressure is directly proportional to the absolute temperature. If an amount of heat is taken up by a gas some of the heat is converted into internal energy and the balance is used in the work done in pushing back the environment as the gas expands.

The gas equation:

For one mole of gas, the preceding laws are combined in the gas equation PV = RT where R = the gas constant = 8.314 joules per Kelvin per mole. The constant for dry air is 2.87 when P is expressed in hectopascals [hPa]. Ordinary gases do not behave exactly in accordance with the gas laws because of molecular attraction and repulsion. The gas equation gives the behaviour of a parcel of air when temperature or pressure, or both, are altered; e.g. if temperature rises and pressure is constant, then volume must increase — consequently the density of the air decreases and the parcel becomes more buoyant. Conversely, if temperature falls and pressure is constant then volume must decrease, the air becomes denser and the parcel less buoyant. Warmed air is comparatively light and cooled air is comparatively heavy. (In meteorological terms a parcel is a mass of air small enough that the whole mass moves or behaves as a single object.)

The equation of state:

P = RrT / M where r = density and M = molecular weight. But for meteorological purposes M is ignored and the equation used is P = RrT. For example, if density remains constant and the temperature increases (decreases), then static pressure increases (decreases) or conversely, if density remains constant and the pressure increases (decreases) then temperature increases (decreases). Or, if pressure remains constant then an increase in temperature causes a decrease in density, and vice versa.

Dalton's law:

The total pressure of a mixture of gases or vapours is equal to the sum of the partial pressures of its components. The partial pressure is the pressure that each component would exert if it existed alone and occupied the same volume as the whole.

Basic atmospheric forces

The basic forces acting in the atmosphere are:

• Gravity, which acts vertically downwards
  
  
• The vertical pressure gradient force, which acts vertically upwards; and the horizontal pressure gradient force, which acts horizontally. Expanded in the following section 1.3.
  
  
• Coriolis effect, which induces turning or curvature in horizontal air flow; refer to section 1.14.2.
  
  
• Friction, which acts to retard horizontal air flow particularly at the surface; refer to sections 3.3.2, 6.3 and 9.1.

1.3 Atmospheric pressure gradient and buoyancy

Atmospheric pressure reflects the average density and thus the weight of the column of air above a given level. Thus the pressure at a point on the earth's surface must be greater than the pressure at any height above it. An increase in surface pressure denotes an increase in mass, not thickness, of the column of air above the surface point. Similarly a decrease in surface pressure denotes a decrease in the mass. The gradient is the difference in pressure vertically and horizontally.

The air throughout the column is compressed by the weight of the atmosphere above it. Thus the density of a column of air is greatest at the surface and decreases exponentially with altitude as shown in the following graph, which is a plot of the rate of decrease in density with increase in altitude. The plot is for dry air at mid-latitudes. ( Mid-latitudes are usually accepted to be the areas between the 30° and 60° parallels, while low latitudes lie between the equator and 30°, and high latitudes between 60° and the poles.) The atmosphere at about 22 000 feet has only 50% of the sea level density. Density decreases by about 3% per 1000 feet between sea level and 18 500 feet, and thereafter the density lapse rate slows.

The dry air density gradient in mid-latitudes. Refer to the table in section 2.3.

As the pressure decreases with height so, in any parcel of air, the downwards pressure over the top of the parcel must be less than the upwards pressure under the bottom. Thus within the parcel there is a vertical component of the pressure gradient force acting upward. Generally this force is balanced by the gravitational force, so the net sum of forces is zero and the parcel floats in equilibrium. This balance of forces is called the hydrostatic balance. When the two forces do not quite balance, the difference is the buoyancy force. This is the upward or downward force exerted on a parcel of air arising from the density difference between the parcel and the surrounding air.

Atmospheric pressure also varies horizontally due to air mass changes associated with the regional thickness of the atmospheric layer. The resultant horizontal pressure gradient force, not being balanced by gravity, forces air to move from regions of higher pressure towards regions of lower pressure. But the movement is modified by the Coriolis effect. The horizontal force is about 1/15 000 of the vertical component.

(Advection is the term used for the transport of momentum, heat, moisture, vorticity or other atmospheric properties, by the horizontal movement of air: refer to section 1.7.5.)

The following graph plots the average mid-latitude vertical pressure gradient and shows how the overall vertical decrease in pressure — the pressure lapse rate — slows exponentially as the air becomes less dense with height. In a denser or colder air mass the pressure reduces at a faster rate. Conversely, in less dense, or warmer, air the pressure reduces at a slower rate. (The hydrostatic equation states that the vertical change in pressure between two levels in any column of air is equal to the weight, per unit area, of the air in the column.) If two air columns have the same pressure change from top to bottom, the denser column will be shorter. Conversely, if the two columns have the same height, the denser column will have a larger change in pressure from top to bottom.

In the ICAO standard atmosphere (details in section 2.3) the rate of altitude change for each 1 hPa (or millibar [mb]) change in pressure is approximately:

The change in altitude for one hectopascal change in pressure can be calculated roughly from the absolute temperature and the pressure at the level using the equation:
altitude change = 96T/P feet.

Atmospheric oxygen and partial pressure

In the homosphere each gas exerts a partial pressure, which is the product of the total atmospheric pressure and the concentration of the gas. As oxygen represents about 21% of the composite gases, the partial pressure of oxygen is about 21% of the atmospheric pressure at any altitude within the homosphere.

Interpolating from the pressure gradient graph above, oxygen partial pressure at selected altitudes is shown below. The decreasing partial pressure of oxygen as an aircraft climbs past 10 000–12 000 feet has critical effects on aircrew; the maximum exposure time — for a fit person — without inspiring supplemental oxygen, is shown in the right-hand column. Perception gradually decreases within the exposure times and exposure beyond these times leads to unconsciousness.

For further information see 'Physiological effects of altitude' in the Flight Theory Guide.

1.4 Atmospheric tides

In the low latitudes a semi-diurnal pressure variation is quite noticeable. Atmospheric pressure peaks at about 1000 hours and 2200 hours local solar time, with minima at 1600 and 0400. The semi-diurnal pressure variation at Cairns in tropical Australia is about 2 hPa either side of the mean; i.e the pressure might be 1015 hPa at 0400, 1019 hPa at 1000, 1015 hPa at 1600 and 1019 hPa at 2200. Meteorologists adjust the daily pressure observations to remove the tide effect.

The atmospheric tide is associated with lunar and solar gravitation, solar heating, and resonance. The tide is not apparent in latitudes greater than 50°–60°. The atmospheric tide is an internal gravity wave with a 12-hour frequency. The semi-diurnal pressure variation is similar to the semi-diurnal gravity variations at the earth's solid surface, the solid earth being subject to tides — the solid tide. A point on the earth's surface moves up and down by as much as 50 cm, with maximum gravity occurring at 1000 and 2200 hours

1.5 Atmospheric moisture

Gas molecules normally exert attractive forces on each other except when in very close proximity, where the interaction is repulsive. If a gas or vapour is cooled so that molecular movements become relatively sluggish, the attractive forces draw the molecules close together to form a liquid. This process is condensation.

A moist atmosphere including water vapour is slightly less dense than a dry atmosphere, at the same temperature and pressure, because the vapour displaces a corresponding amount of the other gases per unit volume and the molecular weight ratio of water vapour to dry air is 0.62:1. Thus a parcel of moister air is more buoyant than surrounding drier air.

Vapour partial pressure is a measure of the amount of water vapour included in a parcel of air and increases as the amount of vapour increases. Moist air — including the maximum amount of water vapour that can be included without condensation at the prevailing temperature — is saturated; i.e. the water vapour pressure is equal to its maximum under that particular condition, and is in equilibrium with a surface of liquid water (e.g. an ocean surface or a water droplet, a water-filled sponge or seasoned wood) at the same temperature. When in equilibrium the same number of water molecules are condensed from the air back into the moist surface as are evaporated from the surface into the air. Water vapour and adjacent moist bodies are always striving for equilibrium, and the equilibrium state is achieved at the saturation vapour partial pressure, the level of which is a function of temperature.

If saturated air is cooled it becomes supersaturated and the excess water vapour immediately condenses onto aerosols (microscopic particles — larger than molecules — of dust, smoke, pollution products and salt small enough to remain suspended in the atmosphere) and forms minute water droplets; refer to section 3.1. The overwhelming majority of aerosols in the upper atmosphere are built up in the cosmic radiation processes and are smaller than the wavelength of light, whereas the larger particles are found near the surface. Those condensation nuclei that have a very high affinity with water — such as salt — are termed hygroscopic particles. (The term hygroscopic describes substances that absorb atmospheric water.) Such nuclei, which originate mainly from sea spray or dust containing salt, help in the initiation of condensation; as it will occur on them well before air is saturated — in the case of sodium chloride it is at 78% relative humidity. If the atmosphere were completely without aerosols, no condensation would occur until extreme supersaturation existed. If droplets or ice crystals already exist, condensation will take place upon them.

Air does not 'hold' water — the maximum amount of vapour that can be present depends on temperature only. A warm atmosphere has greater capacity for water vapour than a cold one; e.g. one kg of air at 35 °C can include 40 grams of water whereas one kg of air at –15 °C can include only one gram. Generally the atmosphere at a tropical ocean surface is 60 times moister than that at 15 000 feet over polar regions.

Saturation vapour partial pressure and dewpoint temperature

The graph above plots the saturation vapour partial pressure, over a liquid water surface, for air temperatures between –20 °C and 45 °C.

The dewpoint is the temperature to which moist air must be cooled, at a given pressure and water vapour content, for it to reach saturation over a water surface. Condensation occurs when the temperature falls below dewpoint; e.g. from the graph above it can be seen that an air parcel at 25 °C and 20 hPa vapour partial pressure would reach its dewpoint on the curve if it were cooled below 17 °C. Very dry air can have a dewpoint well below 0 °C. At ground level if dewpoint is below freezing, a light, crystalline hoar frost forms; but if dew forms before ground temperature subsequently falls below freezing then frozen, or white dew, results.

The frostpoint is the point to which moist air must be cooled for it to reach saturation over an ice surface (e.g. airborne ice crystals). Further cooling induces direct deposition of ice onto solid surfaces, including ice surfaces.

The saturation partial pressure at temperatures below freezing differs for water and ice surfaces. Thus it is possible that air is supersaturated, relative to ice crystals, but unsaturated for supercooled liquid droplets. Refer to section 3.5.2.

The very low saturation partial pressures between –40 °C and –60 °C, corresponding to temperatures at the tropopause, indicate that only minute amounts of water vapour can pass through the tropopause into the stratosphere.

Atmospheric humidity is usually described as a percentage of the saturation value:

Relative humidity [RH]

is the ratio of the amount of water vapour in a parcel of air to the amount that would be present at saturation point, at the same temperature, and is usually expressed as a percentage; i.e. actual density / saturation density x 100. RH can also be calculated as vapour partial pressure / saturation vapour pressure x 100. Thus from the preceding graph, a parcel of air at 25 °C and 20 hPa partial pressure would reach the saturation curve at 32 hPa pressure; therefore the existing relative humidity is 20 / 32 x 100 = 62%.

Note that if the temperature of an air parcel changes then RH changes. For example, during the evening the temperature falls and the RH increases. If 100% RH is reached condensation commences — evening mist. RH does not indicate the actual amount of vapour present, nor the 'humidity' people feel.

Other measures are:

Specific humidity

is the mass of water vapour per unit mass of moist air in grams per kg.

Absolute humidity

is the mass of water vapour per unit volume of air usually in grams per m³.

Humidity mixing ratio

is the ratio of the mass of water vapour to the mass of dry air in a given volume and is usually expressed as grams per kg. It is normally very close to specific humidity except in very humid air.

Wet bulb temperature

is the lowest temperature to which air (surrounding the thermometer bulb) can be cooled by the evaporation of water. The higher the humidity, the slower the evaporation — which ceases at 100% relative humidity when the wet bulb temperature will equal the normal dry bulb temperature.

The spread between surface temperature and dewpoint temperature is an indication of relative humidity and the convection condensation level; e.g. the cloud base may be 1000 feet agl for each 2 °C of spread but inversions, turbulence, etc. will modify this. If the spread is less than 1.5 °C then ceiling and visibility may go below VFR minima. But at 2 °C or greater, CAV may be marginal to OK. Cloud scraps seen to be forming near the surface are a forewarning of visibility problems at low levels.

1.6 Evaporation and latent heat

The amount of moisture contained in the atmosphere at any one time is about 13 000 km³ of water and is equivalent to a world-wide precipitation of 25 mm. As the annual world-wide precipitation is about 850 mm, it follows that the atmospheric moisture is being replenished by evaporation about 35 times per year, or every 10 days or so. About 85% of the moisture evaporates from the oceans, the balance evaporating from fresh-water sources, moist earth and transpiration from plants. Vaporisation is the process of conversion of a substance from the liquid into the vapour state. Fusion is the conversion from solid to liquid state; e.g. snow crystals to rain.

Molecules of water in a condensed state are held to one another by strong forces of attraction, which are balanced by equally strong repulsive forces. Tending to overcome the potential energy of attraction is the escaping tendency of molecules, arising from their kinetic energy. The kinetic energy, and thus the escaping tendency, is a function of absolute temperature. At each temperature a certain fraction of the molecules possesses enough kinetic energy to overcome the forces of attraction of the surrounding molecules and to escape from the surface of the water as vapour — whether that surface is an ocean or a cloud droplet. As the molecules that possess excessive kinetic energy (heat) evaporate from the liquid, the average kinetic energy of the remaining molecules decreases and the temperature drops. The energy carried away with the water vapour, about 2500 joules per gram of vapour, is the latent heat of vaporisation. Conversely, when water vapour condenses back into the liquid state, the latent heat of condensation is released into the surrounding air as sensible heat and has a significant effect on the saturated adiabatic lapse rate. Sensible heat is a function of air temperature while latent heat is a function of H₂O content in its various phases.

Ice melts at 0 °C and requires 330 joules per gram — the latent heat of fusion. If ice is converted directly to water vapour, at the same temperature, it takes about 2800 joules per gram — the latent heat of sublimation. Sublimation is also the process where water vapour is converted directly to ice; e.g. hoar frost forming on a chilled windscreen during take-off or carburettor icing.