Cloud formation and precipitation - Snow - Climate Policy Watcher (2023)

Last Updated on Sun, 06 Aug 2023 | Snow

Moisture in the atmosphere occurs principally in its gaseous phase, as water vapor, but also condenses to form clouds of water droplets or ice crystals. Vapor condenses when its partial pressure exceeds a saturation value, as determined by equilibrium conditions between the vapor and liquid or vapor and ice phases (Section 2.2.2). Since the saturation pressure decreases nearly exponentially with

temperature, warm air at 25 °C can hold about fifty times more vapor than subzero air at -25 °C. Thus, the amount of vapor (measured as the depth of precipitable water vapor in a column extending upwards from the earth's surface) varies widely with season and latitude, from around 1 mm in Arctic continental air in winter to around 60 mm over southern Asia during the monsoon in summer (Barry and Chorley, 2003). The lapse in atmospheric temperature with altitude causes a sharp drop in vapor content and, therefore, about 80% of water vapor is contained within the 1000-700 hPa layer or the lowest 3 km of the atmosphere.

Clouds generally form when warm surface air moves upward, cools adiabatically to saturation, and condenses on condensation or ice nuclei. Precipitation occurs when the cloud particles grow large enough to fall and reach the ground before evaporating or sublimating. Much of the precipitation in high and mid-latitudes begins as snow at higher altitudes and melts to rain as it falls. The occurrence and intensity of precipitation depends on the availability of water vapor and on the concomitant mechanisms for nucleating and growing the particles. Open water in oceans and large lakes forms the major source of moisture in winter. Maximum winter evaporation rates occur when cold continental air blows across warm ocean currents - conditions most often found over the western North Pacific and North Atlantic oceans (Barry and Chorley, 2003). Although snow cover provides an unlimited source of moisture, its evaporation rate is restricted by the low saturation pressures associated with temperatures below 0 °C. Likewise, snowfall amounts in polar regions are relatively limited by the low vapor content of extremely cold air.

Precipitation and clouds are often classified according to the type of vertical air movement leading to their formation. These include (a) frontal or cyclonic lift of air in association with widespread low-pressure systems, (b) orographic lift over topographic barriers, and (c) convective (or buoyant) lift due to heating of the lower atmosphere by land or water surfaces. Table 2.1 (from Schemenauer et al., 1981) lists the typical kinds of snowfall associated with various cloud types. Nimbostratus clouds produce heavier and more persistent snowfalls, while brief but heavy snow showers come from cumulonimbus clouds. When extremely cold conditions prevail at the surface, such as in Antarctica, snow precipitation may occur without significant vertical uplift, just from the condensation of humid air cooled in the vicinity of the surface. Snow crystals then precipitate without any observable clouds. This phenomenon is called "diamond dust."

2.1.2 Snow formation and crystal type

Snow consists of particles of ice that form in the clouds, grow initially by vapor deposition, and then reach the ground before evaporating or melting. This definition

Table 2.1 Types of snowfall associated with various cloud types. (from Schemenauer et al., 1981, with permission from The Blackburn Press).

Front Cloud type

Possible snowfall

Warm Cirrus and derivatives (Ci) Altocumulus (Ac) Altocumulus castellanus (Acc) Altostratus (As)

Nimbostratus (Ns)

Stratocumulus (Sc) Stratus (St)

Cold Cumulus (Cu) and towering cumulus (Tcu) Cumulonimbus (Cb)

Usually virga (snow trails); light snow showers

may occur from the Acc.

Light continuous or intermittent snow from As; however, when the snow is heavy, the As has probably graduated to a nimbostratus cloud.

Continuous snow. Virga occurs from both Ns and As.

Intermittent light powdery snow (fine flakes).

Continuous light powdery snow. (This is the frozen precipitation analogue of warm precipitation drizzle.)

Light snow showers possible; more likely from the Tcu.

Moderate to heavy snow showers.

excludes hail and sleet, which form through the freezing of water, as well as frost, which forms through vapor deposition on the ground. Snow occurs as single crystals or as aggregations of many crystals joined together to form snowflakes. The shape or habit of a single crystal varies from simple hexagonal columns or plates to the complex dendritic or star shapes so favored by artists.

Conditions for snow to form include atmospheric temperatures less than 0 °C and the presence of supercooled water. Although bodies of fresh water (such as lakes) freeze near 0 °C, cloud droplets of pure water can coexist with ice particles down to temperatures as low as -40 °C. Snow begins as ice crystals, which nucleate from the chance aggregation of water molecules into stable ice-like structures called ice embryos. Nucleation can occur either homogeneously or heterogeneously. In the latter process, crystals nucleate onto the surfaces of ice nuclei, which serve to lower the free energy barrier to ice formation. Homogeneous nucleation does not occur in the atmosphere, except perhaps in cirrus clouds, where supercooled drops may freeze spontaneously into cirrus ice (Heymsfield and Miloshevich, 1993; Pruppacher, 1995). Pruppacher and Klett (1997) discuss four mechanisms for heterogeneous nucleation, including the freezing of supercooled water droplets and the direct deposition of water vapor onto ice nuclei.

Major sources of aerosol particles for ice nucleation are dust (commonly silicate minerals of clay) and combustion products from industrial plants, volcanoes, and forest fires. While the concentrations of ice nuclei are generally higher in continental air masses, global measurements show no systematic variation with location

Figure 2.1. Basic hexagonal ice crystal form.

Figure 2.1. Basic hexagonal ice crystal form.

and suggest that nuclei are active far from their origins (Pruppacher and Klett, 1997). Splinters of ice broken from existing snow particles by wind can serve as secondary nuclei, thereby multiplying ice particle concentrations by many orders of magnitude. Unlike condensation nuclei, ice nuclei are generally restricted to substrates similar in structure to ice and are only effective at larger sizes, with diameters typically ranging between 0.1 and 15 |~im. Because the ice-forming process is so selective, the fraction of aerosol particles acting as ice nuclei may be as small as one in 107. Since the critical size of the ice embryo decreases with increasing supersaturation (see Table 2.3 in Section 2.2.2), smaller nuclei become activated at lower temperatures. Based on a review of published data, Fletcher (1962) suggested that the number of ice nuclei increases nearly exponentially with decreasing temperature, from typical concentrations of 0.011-1 at —10 °Cto1001-1 at —30 °C.

Once nucleated, ice crystals grow by mass diffusion of water vapor onto their surfaces in a mechanism called deposition. Ice crystals in a supercooled cloud will grow at the expense of water droplets, because the vapor pressure over ice is less than that over water. Vapor is thus directed from the droplet to the ice surface (see Section 2.2.2). Ice crystals in the early stages of growth are usually less than 75 in diameter and are simple in shape (Schemenauer et al., 1981). The basic shape common to all ice crystals is a hexagonal prism with two basal planes and six prism planes (see Fig. 2.1), which originates because of the covalent bonding of oxygen to hydrogen within the water molecule (Hallett, 1984). The relative growth rates of the basal and prism faces vary with temperature and supersaturation, giving rise to a wide variety of crystal shapes. As indicated by the diagram in Fig. 2.2 (Pruppacher and Klett, 1997, Fig. 2.36: b), crystal habit switches back and forth between plates (radial growth) and columns (axial growth) at transition temperatures of around —3 to —4, —8 to —10, and —20 to —25 °C (Mason, 1971; Frank, 1982; Hallett, 1987; Colbeck et al., 1990). The degree of supersaturation determines secondary crystal features and the crystal growth rate. Higher saturations favor increasingly hollow

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Figure 2.2. Diagram showing the variation of ice crystal shapes with temperature and with excess vapor density (after Pruppacher and Klett, 1997, Fig. 2.36:b, with kind permission from Kluwer Academic Publishers; based on laboratory observations of Kobayashi, 1961, with permission of Taylor & Francis Ltd, http://www.tandf.co.uk/journals; and Rottner and Vali, 1974, copyright 1974 American Meteorological Society).

and skeletal forms, which appear toward the top of the diagram. The difference in vapor pressure between water and ice saturation peaks around -12 °C (see Fig. 2.5 below), contributing to higher growth rates for crystals with dendritic features.

When an ice crystal grows to a size where it has a significant downward velocity and can survive sublimation during its fall to the ground, it becomes a snow crystal. Snow crystals that continue to grow by vapor deposition tend to keep their characteristic shape and proportions. The bulk densities of most crystals are less than that for pure ice, particularly for the more complex or hollow forms. Detailed observations as well as theoretical models have resulted in dimensional and massdiameter relationship for various crystal types. Pruppacher and Klett (1997) provide a useful synopsis of these relationships. Individual unrimed crystals fall at velocities between about 10 and 80 cm s-1 (Kajikawa, 1972), with larger and simpler crystals falling the fastest. Snow crystals reaching the ground typically range by two orders of magnitude in their maximum dimensions from 50 pm to 5 mm. At

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larger dimensions, growth by deposition slows because there is more surface to expand per increment of particle size. Two additional mechanisms, accretion and aggregation, then become important for the growth of snow particles.

Snow crystals exceeding a critical size can grow through collision and coalescence with supercooled cloud drops, which subsequently refreeze, in a process referred to as riming or accretion. The observed width for the onset of riming varies from about 30 for needles to 800 for dendritic crystals. Wang and Ji (2000) successfully use a numerical model to compute the collision efficiencies and threshold diameters of hexagonal plates, broad-branched dendrites, and hexagonal columns. Growth by accretion is relatively fast, with single crystals growing to dimensions of 1-2 mm in 10-20 min (Schemenauer et al., 1981). Because collisions between water droplets and raindrops are less likely to result in coalescence, snow crystals grow larger through accretion than their water-drop counterparts. In order to provide for the dissipation of heat released during freezing, riming typically occurs at subzero temperatures, between —5 and -20 °C. Cases of extreme riming result in the formation of graupel particles or snow pellets. Large graupel particles with dimensions of 3 mm can fall at speeds up to 3 m s—1.

When snow crystals or snow particles collide and "stick" together, snowflakes are formed by the process of aggregation. Snowflakes can consist of between two and several hundred snow crystals. Because of their radiating arms, dendritic crystals aggregate more readily than other crystal types. The aggregation mechanism is most efficient at temperatures near 0 °C, where liquid-like films promote the formation of bonds (sintering) between the particles. Snowflakes therefore have their largest dimension near 0 °C and aggregation is mostly limited to ambient temperatures above —10 °C. Hobbs (1974) predicts that a 1 mm snowflake falling through a cloud of smaller snow crystals can grow to 10 mm diameter in about 20 min. Because of their high air resistance, large snowflakes of up to 15 mm in diameter have fall rates of around 1-2 m s—1.

In summary, snow consists of an intricate variety of snow crystals, as well as rimed and aggregate versions of these forms. The ICSI international classification scheme for snow on the ground (Colbeck et al., 1990) distinguishes eight types of frozen precipitation by shape and growth habitat: columns, needles, plates, stellar dendrites, irregular crystals, graupel, hail, and ice pellets. It is now the standard reference for snow researchers. The older and more elaborate classification scheme of Magono and Lee (1966) is generally supported by recent observations and is recommended for those seeking more detail. Bentley and Humphreys (1931,1962) and Nakaya (1954) provide large collections of fine photographs of these particles. More in-depth material on snow formation can be found in Mason (1971), Hobbs (1974), Schemenauer et al. (1981), Rogers and Yau (1989), and Pruppacher and Klett (1997).

Continue reading here: Thermodynamics of phase equilibria in snow

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