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Aviation Meteorology

The atmosphere and atmospheric thermodynamics


Rev. 22 — page content was last expanded 1 August 2013
  
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1.1 Atmospheric structure

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 molecular temperatures at the 150–300 km level vary between 200 C and 1700 C but due to the rarified atmosphere there is little sensible heat capacity, i.e. a normal thermometer would sense a temperature less than 0 C. The absorbed heat is conducted downward below 100 km where the atmosphere can re-radiate at night.

'Space' is said to start at 100 km altitude, which would mark the thermosphere as the beginning of space.

atmospheric structure diagram

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.

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 ozone (O3) and water vapour (H2O), plus carbon dioxide (CO2), are particularly important as radiation absorbers because of their triatomic structure. The average atmospheric relative molecular mass throughout the homosphere is about 29 atomic mass units (amu).

The relative molecular weight of the main atmospheric components is:

Relative weight of atmospheric gases
(atomic mass units)
Atomic formsDiatomic formsTriatomic forms
H He C O H2 N2 O2 H2O CO2 O3
1 4 12 16 2 28 32 18 44 48

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).There is little or no nitrogen (N2) above 150 km, atomic oxygen (O) dominates between 300 and 1000 km, helium (He) between 1000 and 2000 km, and hydrogen (H) above that.

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 radio transmission waves. But transmissions in the VHF, UHF and higher radio frequencies — which are those used by sport and recreational aviation — 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 older navigation aids such as non-directional beacons which operate in the LF band. 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

Gas laws

The density (the mass of a unit of volume) of dry air is about 1.2 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, so 20 °C equals an absolute temperature of 293 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 static pressure [P] upon the gas; i.e. P × V = constant.

The pressure law:
At a constant volume the static 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 universal gas constant = 8.314 joules per Kelvin per mole. The specific gas constant for dry air (i.e. no water vapour present) 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 approximates 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.)

*Note: 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 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 = 2.87rT. 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. For our purposes pressure is expressed in hectopascals, density in kilograms per cubic metre and temperature in kelvins.

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. As powered recreational aircraft operate at altitudes below 10 000 feet, the component that we have most interest in is the partial pressure of water vapour because that affects the formation of mist, fog and cloud, but above 8000 feet or so the decreasing oxygen partial pressure may start to affect pilot performance — see the 'Physiological effects of altitude'.
Basic atmospheric forces

The basic forces acting in the atmosphere are:



1.3 Atmospheric pressure and buoyancy

The pressure gradient

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.

atmospheric density gradient

The dry air density gradient in mid-latitudes. See the 'International Standard Atmosphere'.


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. A visible application of this principle is readily apparent in the hot-air balloons and airships of sport and recreational aviation.

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, moves air (as wind) from regions of higher pressure towards regions of lower pressure. But the movement is modified by the Coriolis effect. The horizontal force is very small, being 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: see 'Heat advection')

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.

pressure gradient

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

0 to 5000 feet:30 feet/hPa or 34 hPa per 1000 feet
5000 to 10 000 feet:34 feet/hPa or 29 hPa per 1000 feet
10 000 to 20 000 feet:43 feet/hPa or 23 hPa per 1000 feet
20 000 to 40 000 feet:72 feet/hPa or 14 hPa per 1000 feet

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, including water vapour, 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.

Altitude
(ft)
O2
  pressure  
(hPa)
Maximum
exposure
time
  Sea level  210
7000165
10 000150
15 00012030+ minutes
18 00010520–30 minutes
25 000803–5 minutes
30 000651–3 minutes
35 0005030–60 seconds
40 00030  10–20 seconds  

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





1.4 Atmospheric moisture

Water vapour partial pressure, saturation and density

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 and water vapour is the only atmospheric gas that displays this property in nature.

A moist atmosphere that includes 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 slightly more buoyant than surrounding drier air.

*Note: referring to the 'Relative weight of atmospheric gases' table above, the mass of a molecule of H2O is 18 atomic mass units (amu) while that of the oxygen (O2) and nitrogen (N2) diatomic molecules, that make up 21% and 78% of the atmospheric gases, is 32 amu and 28 amu respectively. So the mass of a dry air molecule averages about 29 amu and water vapour mass (18 amu) is 62% of that. As the water vapour molecule occupies about the same space as the dry air molecules it displaces, so air density (mass per unit volume) decreases a little as the humidity of the air increases, and this should be considered when calculating density altitude. Air doesn't 'hold' water, rather the water vapour molecules 'displace' air molecules.

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 occurring 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 the minute water droplets of mist or cloud; see 'Cloud formation'. 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 — substances that absorb water vapour from the air. 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 cloud droplets or ice crystals already exist, condensation will take place upon them.

The maximum amount of vapour that can be present depends on temperature. A warm atmosphere has greater capacity for water vapour than a cold one; e.g. one kg of air at 35 C can include 35 grams of vapour 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.

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

saturation pressure and dew point
Saturation vapour partial pressure and dew point temperature

The dew point is the temperature to which moist air must be cooled, at a given pressure and water vapour content, for it to reach saturation. Condensation occurs when the temperature falls below dew point; 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 dew point on the curve if it were cooled below 17 C. Very dry air can have a dew point well below 0 C. At ground level if dew point 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.

Note: the spread between surface temperature and dew point 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.

The frost point 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 in clouds, but unsaturated for supercooled liquid cloud droplets. See 'Snow'.

Saturation vapour pressure/temperature over ice/water
Ambient temperature C: 0 –10 –20 –30 –40 –60
SVP over water (hPa): 6.1 2.91.3 0.5 0.2-
SVP over ice (hPa): 6.1 2.6 1.0 0.4 0.1 0.01

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.

Quantifying atmospheric humidity
Specific humidity
is the mass of water vapour per unit mass of the moist air in grams per kg.

Humidity mixing ratio
is the ratio of the mass of water vapour to the mass of dry air expressed as grams of vapour per kilogram of dry air. It is normally very close to specific humidity except in very humid air.

Dry bulb temperature
is the ambient or outside air temperature (OAT)— the heat content of the air.

Wet bulb temperature
is the lowest temperature to which the ambient air (flowing around a moistened thermometer bulb) can be cooled by the evaporation of water; see 'Evaporation and latent heat'. The greater the humidity, the lesser the evaporation — which ceases at 100% relative humidity, when the wet bulb temperature will equal the normal dry bulb temperature. If air at 100% relative humidity is cooled, dew point is reached and condensation starts to occur. The dry and wet bulb thermometers comprise a hygrometer; a hygroscope just indicates change in humidity.

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 partial pressure; therefore the existing relative humidity is 20 / 32 x 100 = 62%. The humidity level published in daily weather reports is the relative humidity.

Note that if the temperature of an air parcel changes, then the RH changes. For example, if the amount of water vapour present remains constant, RH decreases when air temperature increases and vice versa. During the evening the temperature falls and the RH increases — if 100% RH is exceeded condensation (evening mist) appears. RH does not indicate the actual amount of vapour present, but in hot weather an increase in RH makes people feel hotter because of the decreased evaporation of perspiration; we rely on evaporative cooling for body temperature control in hot weather.

The table below provides relative humidity if the dry bulb and wet bulb temperatures are known. Airfield altitude has a very slight effect in Australia as there are few airfields/airstrips with an elevation exceeding 4500 feet.

Relative humidity table
Dry bulbDifference between dry bulb and wet-bulb temperatures
Relative humidity
°C-1°-2°-3°-4°-5°-6°-7°-8°-9°-10°
20°91%83%75%67%59%52%45%38%31%26%
25°92%85%77%71%64%57%51%45%39%33%
30°93%86%79%73%67%61%55%50%45%40%
35°93%87%81%75%69%64%59%54%49%44%
40°94%88%82%77%71%66%62%57%52%48%
45°94%89%83%78%73%69%64%60%55%51%
Rule of
thumb value
95%90%85%80%75%70%65%60%55%50%




1.5 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.

latent heat

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 heat energy carried away with the water vapour, about 2500 joules per gram of vapour, is the latent heat of vaporisation (this is the principle of the wet-bulb thermometer). Conversely, when water vapour condenses back into the liquid state, the latent heat of condensation is released into the surrounding air as sensible heat (that increases the air temperature) 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 changing its state, e.g. from liquid to gas.

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.





1.6 Atmospheric and solid Earth 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 Earth tide — caused by lunar/solar gravitation. A point on the Earth's surface might move up and down by as much as 50 cm, with maximum gravitation occurring every 12 hours or so. The solid tide movement is something to be considered in future aircraft GNSS precision approach and landing systems.






Next – atmospheric dynamics The next section of the Aviation Meteorology Guide covers more thermodynamics plus atmospheric dynamics



Aviation meteorology guide modules

| Meteorology guide contents | The atmosphere and thermodynamics (part 1) | Thermodynamics (2) and dynamics |

| Effects of altitude — contained in the Flight Theory Guide module 2 & module 3 |

| Cloud, fog and precipitation | Planetary-scale tropospheric systems | Synoptic scale systems |

| Southern hemisphere winds | Mesoscale systems | Micrometeorology — atmospheric turbulence |

| Airframe and engine icing | Atmospheric electricity | Atmospheric light phenomena |

| Aviation weather reports and forecasts |



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