vertical air motion

vertical air motion In contrast to horizontal air motion (i.e. wind), vertical motion is both difficult to measure and is typically of very small magnitude. As a weather-producing agent, however, it is the key to differences between sunny conditions or cloud and rain. Air may ascend or descend for a variety of reasons. It may be forced to rise and sink over mountains or through convergence or divergence of horizontal airflow. Changes in the horizontal velocity of wind with height (wind shear) may also induce vertical motion. The most significant vertical motion develops under conditions of thermodynamic instability, where semi-isolated parcels of air rise under buoyancy with respect to the surrounding atmosphere. This form of ascent may readily be seen in cumulus clouds, and is termed convection. The degree of uplift from forced ascent is determined by the thermodynamic state of the atmosphere.

Convection in the atmosphere, in its simplest form, results from the atmosphere being heated from the surface upwards. Solar radiation passes through the Earth's atmosphere without raising its temperature. The Earth's surface is thus the first layer to be heated. On average, the temperature in the troposphere decreases with height. The rate of this decrease changes from place to place, and from one moment to the next, as a function of the degree of surface heating, the nature of the surface, the presence of temperature advection aloft, and so on. This rate of temperature decrease, the environmental lapse rate, must be measured to be known; this is done routinely by balloons carrying instruments called radiosondes.

Parcels of air, or thermals, which are isolated from the surrounding atmosphere, cool at a predetermined rate known as the adiabatic lapse rate. Since pressure decreases vertically in the atmosphere, thermals move from a region of higher pressure near the surface to lower pressure, a change which allows them to expand. Expansion means that work is being done on the surrounding atmosphere. That work is paid for by a loss of internal energy, or temperature. As the temperature is lowered, so the ability of the air to hold water in gaseous form decreases. If a critical temperature, known as the dew point, is reached, condensation occurs and clouds form. Similarly, air which is sinking contracts. The atmosphere is doing work on the parcel of air, which gains internal energy, expressed by a temperature increase. Water evaporates back into vapour form and the cloud disappears. This rate of change of temperature with height is fixed at 10 °C km⁻¹ in dry air and is known as the dry adiabatic lapse rate Following condensation (i.e. above cloud base), latent heat of condensation is released into the parcel of air. Latent heat is released because water changes state from a gaseous form (water vapour) to a liquid form (cloud water droplets), which corresponds to a change to a lower energy level. This release of latent heat acts to slow the rate of cooling from 10 °C km⁻¹ to an average of 6 °C km⁻¹, a rate of change of temperature known as the saturated adiabatic lapse rate. It is important to point out that in the case of saturated air (above cloud base) this rate of change of temperature with height is not fixed, since it depends on the amount of water vapour initially in the air and on the temperature of the air itself. A warm, moist air mass will cool very slowly after condensation, since the release of latent heat following condensation will be large. A cool, dry air mass, if lifted to condensation level, will cool at a rate not much slower than that of dry air. In other words, in cool, dry air the adiabatic lapse rate and the saturated adiabatic lapse rate are closer together.

The stability or instability of an air mass depends on whether the isolated rising parcel (thermal), which cools at a fixed rate determined by the dry adiabatic lapse rate to condensation level, and by the saturated lapse rate thereafter, finds itself at any moment to be warmer or cooler than the surrounding air. If warmer, the parcel is less dense and buoyancy forces will force it to rise further. The air is said to be unstable. If cooler, it will be more dense and will sink. The air is said to be stable. When air is stable, there is no convection: air which is displaced up or down by mountains, for example, will return to its previous altitude. Unstable air is associated with convection.

These relationships are shown graphically in Fig. 1a, the environmental lapse rate, or measured lapse rate, being represented by the line AB. Temperature decreases rapidly from the surface. This may result for a variety of reasons. Cold air may have been advected aloft, or the surface may have been intensely heated. The profile of line AB will change with time. Parcels of air rising from the surface will cool at a predetermined rates, given by the line AC. In this case we assume that no condensation has occurred and that cooling is 10 ^oC km⁻¹ (dry adiabatic lapse rate). We shall also not concern ourselves with what causes the parcel to rise just yet. We can determine whether the atmosphere is unstable or stable by comparing the temperature of the ascending parcel (line AC) with that of the surrounding atmosphere (line AB) at any common altitude. At all altitudes, the temperature of the parcel exceeds that of the surrounding atmosphere. The parcel rises by its own buoyancy.

Figure 1b illustrates a stable case. Here, the environmental lapse rate, AB, increases with height, a condition known as a temperature inversion. Forced parcel ascent will occur along line AC, the dry adiabatic lapse rate. Any ascent by the parcel will lead to the parcel becoming cooler and more dense than its surroundings. It will therefore sink back to its starting point.

A further type of instability needs to be considered, that of conditional instability. Here, rising air may become unstable if there is enough moisture for condensation to occur, or rather if the air is sufficiently close to saturation for condensation to occur. Condensation leads to latent heat release, which further warms the parcel of air and may render it less dense than its surroundings. This is illustrated in Fig. 2a. Line AB represents the environmental lapse rate. Ascent by the parcel occurs at two rates. Initially the parcel cools at the rate shown by AC, the dry adiabatic lapse rate. Condensation of the moist air occurs at point C and further ascent (CD) takes place at a slower rate of cooling, the saturated adiabatic lapse rate. Above point E the parcel becomes warmer than the environment. The atmosphere is classed as conditionally unstable. Figure 2b shows an identical environmental lapse rate. The difference here is that the air is extremely dry. Ascent by the parcel at the dry adiabatic lapse rate is given by line AC. The dry air is unable to cool to condensation level (dew-point temperature) throughout the ascent. Since the dry adiabatic lapse rate is a faster rate of cooling, the parcel remains cooler than the surrounding atmosphere and the atmosphere is stable. Conditional instability therefore depends on the amount of moisture in the parcel.

Ascent forced by mountains, convergence in a low-pressure system, or even wind shear may thus be enhanced if the atmosphere is in an unstable state. Typically, forced ascent is needed to trigger conditional instability. Evidence of this is the tendency for thunderstorms to develop near mountains or in low-pressure systems. The forced ascent provides the energy to raise the air to the altitude where instability is released. Subsequent energy release occurs as a result of the differences between the parcel and its surroundings.

An assumption thus far is that rising air is isolated from the environment. In reality, mixing, known as entrainment, between the parcel and the environment does occur, but is less prevalent than would be expected. Clouds, for example, tend to act as conduits for moist, warm surface air to reach higher levels with minimal mixing.

Differences between rising and sinking air are worth noting. Ascent is often concentrated inside cloud turrets, with values reaching 30ms⁻¹ in vigorous deep thunderstorms. Descent is often very slow, of the order of a few centimetres per hour, and tends to occur over much larger areas. Evidence of slow, large-scale descent is provided in the desert areas of the subtropics, such as the Sahara, which cover vast expanses of the globe.

R. Washington

Bibliography

Meteorological Office (1991) Meteorological glossary. HMSO Publications, London.