Values Of The Evaporation Rate

The global average of the evaporation rate must equal the annual mean rainfall of about 1,020

Figure 4.10 The change of the actual rate of evaporation Ea from a crop as the soil dries out, showing it as a two-stage process. When the soil is initially fully wetted, Ea is the potential rate E,, which depends primarily on atmospheric conditions (i.e. net radiation, temperature, humidity and wind speed), and secondarily on the crop's roughness and albedo. In the first stage of drying, Et is maintained so long as the soil can provide water to the crop at that rate. The water content of the soil M is eventually insufficient to allow evaporation to continue at that rate, and so Ea declines during the second stage. Various measurements have shown that the rate at which soil can supply water is proportional to M2, and Ea falls correspondingly. In this second stage, Ea is limited by the soil's ability to supply water, whereas the limitation in the first stage was atmospheric demand.

Figure 4.10 The change of the actual rate of evaporation Ea from a crop as the soil dries out, showing it as a two-stage process. When the soil is initially fully wetted, Ea is the potential rate E,, which depends primarily on atmospheric conditions (i.e. net radiation, temperature, humidity and wind speed), and secondarily on the crop's roughness and albedo. In the first stage of drying, Et is maintained so long as the soil can provide water to the crop at that rate. The water content of the soil M is eventually insufficient to allow evaporation to continue at that rate, and so Ea declines during the second stage. Various measurements have shown that the rate at which soil can supply water is proportional to M2, and Ea falls correspondingly. In this second stage, Ea is limited by the soil's ability to supply water, whereas the limitation in the first stage was atmospheric demand.

mm (Chapter 10), since the input into the atmosphere must equal its output. This implies an average latent-heat flux from the surface of 80 W/m2, which is a substantial fraction of the net radiation available at the surface (Figure 2.17).

The evaporation rate varies with latitude for reasons associated with the Dalton evaporation equation (Note 4.E). Firstly, the surface water's vapour pressure is affected by its temperature, which varies with latitude (Chapter 11). Table 4.2 shows the importance of evaporation from the oceans, and from the lowest latitudes, where global radiation is greatest (Figure 2.11). Secondly, the air's humidity varies between the pole and the equator (Chapter 6), affecting the

Table 4.2 Effect of latitude on the actual annual evaporation, in millimetres depth per annum

Oceans

Land

Overall

Latitude (°S)

(mm/a)

(mm/a)

(mm/a)

0-10

1,380

1,330

1,370

10-20

1,550

1,050

1,510

20-30

1,390

410

1,310

30-40

1,200

520

1,180

40-50

880

530

860

50-60

560

-

550

Figure 4.11 Variation over the Earth of the annual mean latent-heat flux (W/m2). Multiply the numbers by 12.7 to convert from latent heat flux (W/m2) to annual evaporation (mm/a).

vapour pressure e in Dalton's equation. Thirdly, there are differences of wind strength around the globe (Chapter 12). Fourthly, evaporation inland is less because water is scarcer (Section

The global pattern of actual evaporation rates is illustrated in more detail in Figure 4.11. Values exceed 150 W/m2 (i.e. 1,920 mm/a or 5.3 mm/ day) over subtropical oceans in the southern hemisphere, especially just north of the Tropic of Capricorn where cloudiness is least (Chapter 8). Other high values are seen just east of the continents, locally exceeding the energy available from net radiation—a paradox which is considered in the next chapter. The lowest evaporation values occur at the Poles and over subtropical deserts, such as the interior of Australia, where they are less than 200 mm/a. The average evaporation rate from all the Earth's land is about 420 mm/a, and from the oceans about 1,260 mm/a.

The distribution of pan evaporation Ep is almost opposite to that of actual evaporation Ea. Over the oceans the two are similar, but Ep is more in subtropical deserts, whereas Ea is much less. Ea in a desert might be only 5 per cent of Ep, for instance. Over Australia, Ep ranges from less than 800 mm/a in western Tasmania (i.e. 2.2 mm/d) to more than 16 mm/d in January in central Australia (Figure 4.12).

Rapid Evaporation

Evaporation is generally limited by the available amount of net radiation (Chapter 5), except that the rate can be substantially enhanced in small and isolated wet areas, like an irrigated field in dry country. This happens in two ways. Firstly, a gradually diminishing oasis effect (or edge effect) extends some 50 m into the field from the upwind boundary due to heat and

Figure 4.12 Class-A pan evaporation (mm/d) at places in Australia, in January and July.

dryness in the wind from relatively arid surroundings. Thus, evaporation of 5.3 mm was measured from a plot of irrigated lucerne in a dry area at Deniliquin in New South Wales during 7.5 daylight hours, which is fast, and a field of irrigated lucerne in Nevada evaporated 14.2 mm in one day. The same oasis effect has been found to lead to pan evaporation of up to 24 mm/d within dry surroundings in California.

Secondly, there can be a clothes-line effect in a tall crop or forest due to air flowing through the canopy and not just over it. The throughflow creates extra evaporation from within the canopy.

We have seen that evaporation rates can vary considerably, and hence the fraction of the incoming energy that is consumed as latent heat. This governs how much of the incoming radiant heat remains to warm the surroundings. Reconciliation of these aspects of climate is the subject of the 'energy balance' considered in the next chapter.

4.7 DEW

The opposite of evaporation is condensation (Figure 4.1), involving the same amounts of latent heat but released as sensible heat, instead of being absorbed (Section 4.1; Chapter 5). Some of the atmosphere's water vapour may condense onto surfaces chilled at night by the net loss of longwave radiation (Section 2.7). This is dewfall. It occurs once the air's temperature is such that the corresponding saturation vapour pressure es is less than the actual vapour pressure. It is as though a sponge has been shrunk by cooling so that its capacity becomes less than the volume of fluid it contains. We say the air has become 'supersaturated'. The excess spills out, condensed into liquid.

The condensation of that excess forms fog (Chapter 8) if the air is humid and the nocturnal cooling rapid. Otherwise it forms as dew on nearby surfaces, such as the leaves of plants. These tend to be colder than screen temperature at night (Section 3.5), so the adjacent air is first to become supersaturated.

As much as 0.3 mm of dew may be deposited during a twelve-hour night, liberating 17 W/m2. However, less than that is normally deposited, even though it may seem more when concentrated on the tips of leaves. For instance, the average nightly deposit when dew fell was

0.13 mm during a series of measurements in Sydney There was dewfall only when the air was moist and calm, and the nocturnal sky was clear.

Even the small amount of moisture in dew can be useful in arid areas such as northern Chile. There it is collected by means of piles of stones, where the condensation drips to the inner base of the pile and is then shielded from the following day's sunshine. Dew may also be collected on sloping metal sheets. It makes plants grow at the foot of telegraph poles in the desert.

The latent heat which is released when dew condenses tends to offset the usual cooling at night, so that dewfall governs the subsequent day's minimum temperature (Section 3.4). In fact, the minimum cannot be much below the dewpoint, the temperature at which dew first forms (Chapter 6), because cooling below that causes further condensation which releases compensating latent heat.

Dew does not arise only from the cooling and supersaturation of the adjacent air. An additional kind is caused by distillation from the ground. The ground at night is warmer than the air and may be wet, so that it evaporates moisture into an atmosphere already almost saturated. Condensation follows onto surfaces like the leaves of plants, which cool more rapidly than the ground because of a smaller heat content and wider exposure to the sky.

A third source of moisture on leaves at daybreak is guttation. This is the exudation of moisture from the leaves of some plants, as a carry-over from the transpiration of the previous daytime. Neither guttation nor distillation affect the transfer of heat from the surface to the atmosphere, so they are irrelevant to the flows of energy which are related together in the next chapter.

NOTES

4.A Water molecules

4.B Protection of crops from frost

4.C Saturation vapour pressure and temperature

4.D Rates of evaporation

4.E Dalton's evaporation equation

4.F Effect of drop radius on its evaporation rate

4.G Crop evaporation and yield

4.H The Relative Strain Index of comfort

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