From Precipitation Lands On Lands

SURFACE RUNOFF

TO OCEAN

OCEANS 97% OF ALL WATER

RIVERS 0.01% SOIL MOISTURE 0.05%

LAKES AND MARSH 0.3%

ICE SHEETS. GLACIERS. AND GROUND ICE 70% OF ALL FRESH WATER

GROUNDWATER (<750m) 14%

GROUND WATER (750-4.000 M) 16%

OCEANS (23.4 XI06 km3)

CONTINENTS (Percentages refer to fresh water total)

Figure 4.1 The hydrological cycle and water storage of the globe. The exchanges in the cycle are referred to 100 units, which equal the mean annual global precipitation of 1130 mm. The percentage storage figures for atmospheric and continental water are percentages of all fresh water. The saline ocean waters make up 97 per cent of all water. The horizontal advection of water vapour indicates the net transfer.

Figure 4.1 The hydrological cycle and water storage of the globe. The exchanges in the cycle are referred to 100 units, which equal the mean annual global precipitation of 1130 mm. The percentage storage figures for atmospheric and continental water are percentages of all fresh water. The saline ocean waters make up 97 per cent of all water. The horizontal advection of water vapour indicates the net transfer.

Source: From More (1967) updated after Korzun (1978).

where AQ is the time change of moisture in an atmospheric column, E = evaporation, P = precipitation, Dq = moisture divergence out of the column, AS = surface storage of water, and r = runoff. For short-term processes, the water balance of the atmosphere may be assumed to be in equilibrium; however, over periods of tens of years, global warming may increase its water storage capacity.

Because of its large heat capacity, the global occurrence and transport of water is closely linked to global energy. Atmospheric water vapour is responsible for the bulk of total global energy lost into space by infra-red radiation. Over 75 per cent of the energy input from the surface into the atmosphere is a result of the liberation of latent heat by condensation and, principally, the production of rainfall.

The average storage of water vapour in the atmosphere (Table 4.1), termed the precipitable water content (about 25 mm), is sufficient for only ten days' supply of

Table 4.1 Mean water content of the atmosphere (in mm of rainfall equivalent).

January

July

Northern hemisphere

19 34

Southern hemisphere

25 20

World

22 27

Source: After Sutcliffe (1956).

rainfall over the earth as a whole. However, intense (horizontal) influx of moisture into the air over a given region makes possible short-term rainfall totals greatly in excess of 30 mm. The phenomenal record total of 1870 mm fell on the island of Réunion, off Madagascar, during twenty-four hours in March 1952, and much greater intensities have been observed over shorter periods (see E.2a, this chapter).

B HUMIDITY

1 Moisture content

Atmospheric moisture comprises water vapour, and water droplets and ice crystals in clouds. Moisture content is determined by local evaporation, air temperature and the horizontal atmospheric transport of moisture. Cloud water, on average, amounts to only 4 per cent of atmospheric moisture. The moisture content of the atmosphere can be expressed in several ways, apart from the vapour pressure (p. 24), depending on which aspect the user wishes to emphasize. The total mass of water in a given volume of air (i.e. the density of the water vapour) is one such measure. This is termed the absolute humidity (rw) and is measured in grams per cubic metre (g m-3). Volumetric measurements are seldom used in meteorology and more convenient is the mass mixing ratio (x). This is the mass of water vapour in grams per kilogram of dry air. For most practical purposes, the specific humidity (q) is identical, being the mass of vapour per kilogram of air, including its moisture.

More than 50 per cent of atmospheric moisture content is below 850 mb (approximately 1450 m) and more than 90 per cent below 500 mb (5575 m). Figure 4.2 illustrates typical vertical distributions in spring in middle latitudes. It is also apparent that the seasonal effect is most marked in the lowest 3000 m (i.e. below 700 mb). Air temperature sets an upper limit to water mb 400-|

500-

700-

850-

950-Surface

500-

700-

850-

950-Surface

Figure 4.2 The vertical variation of atmospheric vapour content (g/kg) at Tucson, AZ and Miami, FL at 12 UTC on 27 March 2002.

vapour pressure - the saturation value (i.e. 100 per cent relative humidity); consequently we may expect the distribution of mean vapour content to reflect this control. In January, minimum values of 1-2 mm (equivalent depth of water) occur in northern continental interiors and high latitudes, with secondary minima of 5-10 mm in tropical desert areas, where there is subsiding air (Figure 4.3). Maximum vapour contents of 50-60 mm are over southern Asia during the summer monsoon and over equatorial latitudes of Africa and South America.

Another important measure is relative humidity (r), which expresses the actual moisture content of a sample of air as a percentage of that contained in the same volume of saturated air at the same temperature. The relative humidity is defined with reference to the mixing ratio, but it can be determined approximately in several ways:

where the subscript s refers to the respective saturation values at the same temperature; e denotes vapour pressure.

A further index of humidity is the dew-point temperature. This is the temperature at which saturation occurs if air is cooled at constant pressure without addition or removal of vapour. When the air temperature and dew point are equal the relative humidity is 100 per cent, and it is evident that relative humidity can also be determined from:

es at dew-point e at air temperature

X 100

Figure 4.2 The vertical variation of atmospheric vapour content (g/kg) at Tucson, AZ and Miami, FL at 12 UTC on 27 March 2002.

The relative humidity of a parcel of air will change if either its temperature or its mixing ratio is changed. In general, the relative humidity varies inversely with temperature during the day, tending to be lower in the early afternoon and higher at night.

Atmospheric moisture can be measured by at least five types of instrument. For routine measurements the wet-bulb thermometer is installed in a louvred instrument shelter (Stevenson screen). The bulb of the standard thermometer is wrapped in muslin, which is kept moist by a wick from a reservoir of pure water. The s s s

Figure 4.3 Mean atmospheric water vapour content in January and July 1970 to 1999, in mm of precipitable water.

Source: Climate Diagnostics Center, CIRES-NOAA, Boulder, CO.

Figure 4.3 Mean atmospheric water vapour content in January and July 1970 to 1999, in mm of precipitable water.

Source: Climate Diagnostics Center, CIRES-NOAA, Boulder, CO.

evaporative cooling of this wet bulb gives a reading that can be used in conjunction with a simultaneous dry-bulb temperature reading to calculate the dew-point temperature. A similar portable device - the aspirated psychrometer - uses a forced flow of air at a fixed rate over the dry and wet bulbs. A sophisticated instrument for determining the dew-point, based on a different principle, is the dew-point hygrometer. This detects when condensation first occurs on a cooled surface. Two other types of instrument are used to determine relative humidity. The hygrograph utilizes the expansion/ contraction of a bundle of human hair, in response to humidity, to register relative humidity continuously by a mechanical coupling to a pen arm marking on a rotating drum. It has an accuracy of ±5 to 10 per cent.

For upper air measurements, a lithium chloride element detects changes in electrical resistance to vapour pressure differences. Relative humidity changes are accurate within ± 3 per cent.

2 Moisture transport

The atmosphere transports moisture horizontally as well as vertically. Figure 4.1 shows a net transport from oceans to land areas. Moisture must also be transported meridionally (south-north) in order to maintain the required moisture balance at a given latitude (i.e. evaporation - precipitation = net horizontal transport of moisture into an air column). Comparison of annual average precipitation and evaporation totals

for latitude zones shows that in low and middle latitudes P > E, whereas in the subtropics P < E (Figure 4.4A). These regional imbalances are maintained by net moisture transport into (convergence) and out of (divergence) the respective zones (DQ, where divergence is positive): E - P = dq

A prominent feature is the equatorward transport into low latitudes and the poleward transport in middle latitudes (Figure 4B). Atmospheric moisture is transported by the global westerly wind systems of middle latitudes towards higher latitudes and by the easterly trade wind systems towards the equatorial region (see Chapter 7). There is also significant exchange of moisture between the hemispheres. During June to August there is a moisture transport northward across the equator of 18.8 X 108 kg s-1; during December to February the southward transport is 13.6 X 108 kg s-1. The net annual south to north transport is 3.2 X 108 kg s-1, giving an annual excess of net precipitation in the northern hemisphere of 39 mm. This is returned by terrestrial runoff into the oceans.

It is important to stress that local evaporation is, in general, not the major source of local precipitation. For example, 32 per cent of the summer season precipitation over the Mississippi River basin and between 25 and 35 per cent of that over the Amazon basin is of 'local' origin, the remainder being transported into these basins by moisture advection. Even when moisture is available in the atmosphere over a region, only a small portion of it is usually precipitated. This depends on the efficiency of the condensation and precipitation mechanisms, both microphysical and large scale.

Using atmospheric sounding data on winds and moisture content, global patterns of average water vapour flux divergence (i.e. E - P > 0) or convergence (i.e. E - P < 0) can be determined. The distribution of atmospheric moisture 'sources' (i.e. P < E) and 'sinks' (i.e. P > E) form an important basis for understanding global climates. Strong divergence (outflow) of moisture occurs over the northern Indian Ocean in summer, providing moisture for the monsoon. Subtropical divergence zones are associated with the high-pressure areas. The oceanic subtropical highs are evaporation sources; divergence over land may reflect underground water supply or may be artefacts of sparse data.

C EVAPORATION

Evaporation (including transpiration from plants) provides the moisture input into the atmosphere; the oceans provide 87 per cent and the continents 13 per cent.

The highest annual values (1500 mm), averaged zonally around the globe, occur over the tropical oceans, associated with trade wind belts, and over equatorial land areas in response to high solar radiation receipts and luxuriant vegetation growth (Figure 4.5A). The larger oceanic evaporative losses in winter, for each hemisphere (Figure 4.5B), represent the effect of outflows of cold continental air over warm ocean currents in the western North Pacific and North Atlantic (Figure 4.6) and stronger trade winds in the cold season of the southern hemisphere.

Evaporation requires an energy source at a surface that is supplied with moisture; the vapour pressure in the air must be below the saturated value (es); and air motion removes the moisture transferred into the surface layer of air. As illustrated in Figure 2.14, the saturation vapour pressure increases with temperature. The change in state from liquid to vapour requires energy to be expended in overcoming the intermolecular attractions of the water particles. This energy is often acquired by the removal of heat from the immediate surroundings, causing an apparent heat loss (latent heat), as discussed on p. 55, and a consequent drop in temperature. The latent heat of vaporization needed to evaporate 1 kg of water at 0°C is 2.5 X 106 J. Conversely, condensation releases this heat, and the temperature of an airmass in which condensation is occurring is increased as the water vapour reverts to the liquid state.

The diurnal range of temperature can be moderated by humid air, when evaporation takes place during the day and condensation at night. The relationship of saturation vapour pressure to temperature (Figure 2.14) means that evaporation processes limit low latitude ocean surface temperature (i.e. where evaporation is at a maximum) to values of about 30°C. This plays an important role in regulating the temperature of ocean surfaces and overlying air in the tropics.

The rate of evaporation depends on a number of factors, the two most important of which are the difference between the saturation vapour pressure at the water surface and the vapour pressure of the air, and the existence of a continual supply of energy to the surface. Wind velocity also affects the evaporation rate, because

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