Vertical Convection In The Atmosphere

Processes occurring at the air-sea interface are greatly affected by the degree of turbulent convection that can occur in the atmosphere above the sea-surface. This in turn is dependent on the degree of stability of the air, i.e. on the extent to which, once displaced upwards, it tends to continue rising.

Two ways in which density may vary with height in a fluid are illustrated in Figure 2.9. Situation (a), in which density increases with height, is unstable, and fluid higher up will tend to sink and fluid lower down will tend to rise. Situation (b), in which density decreases with height, is stable: a parcel of fluid (at, say, position O) that is displaced upwards will be denser than its surroundings and will sink back to its original position.

Figure 2.9 Possible variations of fluid density with height, leading to (a) unstable and (b) stable conditions.

Figure 2.9 Possible variations of fluid density with height, leading to (a) unstable and (b) stable conditions.

density increasing stable density increasing

density increasing

The density of air depends on its pressure and its temperature. It also depends on the amount of water vapour it contains - water vapour is less dense than air - but for most practical purposes water vapour content has a negligible effect on density. Thus, the variation of density with height in a column of air is determined by the variation in temperature with height.

The variation in temperature with height in the atmosphere is complicated. For one thing, air, like all fluids, is compressible. When a fluid is compressed, the internal energy it possesses per unit volume by virtue of the motions of its constituent atoms, and which determines its temperature, is increased. Conversely, when a fluid expands, its internal energy decreases. Thus, a fluid heats up when compressed (a well-known example of this is the air in a bicycle pump), and cools when it expands. Changes in temperature that result from changes in volume/density and internal energy, and not because of gain or loss of heat from or to the surroundings, are described as adiabatic. Adiabatic temperature changes have a much greater effect on the behaviour of air masses than do other mechanisms for gaining or losing energy (absorbing or emitting radiation, or mixing with other air masses).

Imagine a parcel of air above a warm sea-surface moving upwards in random, turbulent eddying motions. As it rises, it is subjected to decreasing atmospheric pressure and so expands and becomes less dense; this results in an adiabatic decrease in temperature, which for dry air is 9.8 °C per kilometre increase in altitude (black line on Figure 2.10). If the adiabatic decrease in temperature of a rising parcel of air is less than the local decrease of temperature with height in the atmosphere, the rising parcel of air will be warmer than the surrounding air and will continue to rise. In other words, this is an unstable situation, conducive to upward convection of air. If, on the other hand, the adiabatic cooling of the rising parcel of air is sufficient to reduce its temperature to a value below that of the surrounding air, it will sink back to its original level - i.e. conditions will be stable.

However, the situation is further complicated by the presence of water vapour in the air. Water vapour has a high latent heat content, with the result that the constant 'dry' adiabatic lapse rate of 9.8 °C per km is of limited relevance. If rising air is saturated with water vapour - or becomes saturated as a result of adiabatic cooling - its continued rise and associated adiabatic cooling results in the condensation of water vapour (onto atmospheric nuclei, such as salt or dust particles), to form water droplets. Condensation releases latent heat of evaporation, which partly offsets the adiabatic cooling, so the rate at which air containing water vapour cools on rising (blue 'saturated' lapse rate in Figure 2.10(a)) is less than the rate for dry air.

always stable

always unstable temperature (°C)

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