Precipitation formation

Precipitation is the moisture flux from the atmosphere to the Earth's surface coupling the atmospheric and terrestrial branches of the hydrologic cycle and serving as the downward directed mass flux for climate of the second kind. Rain and snow provide the primary moisture inputs to the land phase of the hydrologic cycle. Dew and fog-drip provide small moisture inputs for specific locations that are very important locally, but the moisture quantity delivered by these processes is small compared to rain and snow. Accurate measurement of rain and snow is essential for hydroclimatic analyses, but precipitation is highly variable in both time and space compared to temperature and pressure. The general problem of data representativeness is particularly challenging regarding precipitation due to its time and space variability which is a function of the processes responsible for producing precipitation.

Precipitation is produced in a series of stages beginning with supersaturation of ascending and cooling air. Condensation, droplet formation, and successful descent of the droplets or particles to the Earth's surface complete the atmospheric cycling of moisture. The amount of water vapor present in the air is a fundamental factor to precipitation formation. Quantifying the amount of water vapor in the atmosphere is achieved by determining the contribution water vapor makes to total atmospheric pressure. This is the partial pressure due to water vapor or the vapor pressure, and it is typically about 2.5 hPa out of the total atmospheric pressure (Trenberth, 1998). The common convention is to view the process as occurring in a defined volume of air identified as an air parcel. Saturation is the state achieved when evaporation and condensation of water molecules in the parcel are in equilibrium. Saturation is quantified by relative humidity (see Equation 3.4) of 100%.

The saturation vapor pressure is the pressure at which the air parcel is saturated with water vapor, and it varies only with temperature (see Fig. 3.4). This relationship has important consequences relative to the atmosphere and precipitation formation. The first feature is that the saturation vapor pressure of warm air is greater than the saturation vapor pressure of cold air when the quantity of water vapor in the volume is unchanged. The second characteristic is that cooling an air parcel is an efficient means for achieving saturation of air as an initial step in precipitation formation.

A key mechanism for cooling air is to induce vertical displacement of the parcel by lifting the parcel to higher elevations. The vertical ascent of the parcel

Lifting and cooling

Lifting and cooling

Vertical Lifting Diagram

Fig. 4.1. Simplified diagram of the precipitation formation process.


Fig. 4.1. Simplified diagram of the precipitation formation process.

results in cooling as the parcel uses internal energy to expand as it moves into a decreasing free atmospheric pressure environment with increasing altitude. The processes resulting in precipitation are depicted in Figure 4.1. The parcel cools initially at about 9.8 °Ckm_1 of ascent. This cooling is known as the dry adiabatic lapse rate. The free atmosphere cools at a mean vertical rate for the troposphere of 6.5 °Ckm_1. However, the tropospheric lapse rate varies with time and space.

Three types of lifting mechanisms are recognized as being responsible for triggering vertical atmospheric motion. Convective lifting is related to turbulent air currents initiated by frictional drag due to rough surface features, eddy circulations in the surface air flow, or thermal differences in surface materials that result in differential heating of the air. Frontal lifting occurs due to the contrasting air mass characteristics along the boundary separating expanses of air with different densities and temperatures. The nature of the advancing air mass determines the frontal slope and the horizontal and vertical character of the lifting. Orographic lifting is associated with the forced ascent of air resulting from horizontal flows encountering a terrain barrier such as mountains. The mountains' height influences the character of the lifting, but wind direction relative to the mountains' orientation influences the overall character of the atmospheric lifting.

The lifting condensation level (LCL) defines the height to which a moist air parcel must ascend for adiabatic cooling to produce saturation of the parcel with respect to a plane surface of water. Once saturation of rising air is achieved, condensation can occur if the air is cooled further. Water vapor requires a surface on which to condense because the process would require rarely observed high degrees of supersaturation of air to allow water molecules to collide and stick to produce a droplet. In the free atmosphere, the condensation surface is provided by microscopic impurities suspended in the air and known as cloud condensation nuclei (CCN). These particles have diameters of 0.1 to

Table 4.1. Selected precipitation types and representative physical characteristics

Precipitation type Intensity (cm h 1) Median diameter (mm) Fall velocity (m s 1)

Fog Snow Light rain Heavy rain Cloudburst


0.003 1

10 mm and originate from natural and anthropogenic sources. The droplets that form around CCN are relatively small and remain buoyant due to the continuous rising air stream from the surface. These droplets may become dense enough to be visible as clouds. Growth of the cloud droplets is necessary for them to achieve a mass with a terminal velocity great enough to exceed the buoyancy provided by the vertical air currents. Droplet growth depends on the cloud temperature where the growth occurs.

In clouds with temperatures above freezing, droplets can grow by collision and coalescence with other droplets. This growth process is enhanced when droplets of different sizes are involved due to differences in terminal velocities. The greater terminal velocities of larger droplets permit them to overtake smaller droplets. Clouds extending above the freezing level support droplet growth by a three-phase process or the Bergeron process. These clouds contain ice and supercooled water above the freezing level. Since the saturation vapor pressure over ice is less than the saturation vapor pressure over water at the same temperature (see Fig. 3.4), water molecules move from supercooled drops to ice crystals. The ice crystals grow while the water droplets become smaller. The terminal velocity of the ice crystals increases, and they fall toward the Earth's surface. The ice crystals melt and arrive as a water droplet if they encounter air above freezing before they reach the surface. Selected characteristics of different precipitation types are listed in Table 4.1.

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