Jpl pg

Allowing for the conversion of the mass of water (mw) to equivalent depth over a unit cross-sectional area, the precipitable water is given by:

0.1 fp2

W (mm) =- SHdp g hi where p is in mb, SH in gkg-1 and g = 9.81 ms-2.

In practice, the integration cannot be performed since q is not known as a function of p. A value of W is obtained by summing the contributions for a sequence of layers in the troposphere from a series of measurements of the specific humidity q at different heights and using the average specific humidity q over each layer with the appropriate pressure difference: 0.1 p2

Example. From a radiosonde (balloon) ascent, the pairs of measurements of pressure and specific humidity shown in Table 1.4 were obtained. The precipitable water in a column of air up to the 250mb level is calculated (g = 9.81 ms-2).

1.3.2 Solar radiation

The main source of energy at the Earth's surface is radiant energy from the Sun, termed solar radiation or insolation. It is the solar radiation impinging on the Earth that fuels the heat engine driving the hydrological cycle. The amount of radiant energy received at any point on the Earth's surface (assuming no atmosphere) is governed by the following well-defined factors.

(a) The solar output. The Sun, a globe of incandescent matter, has a gaseous outer layer about 320 km thick and transmits light and other radiations towards the Earth from a distance of 145 million km. The rate of emission of energy is shown in Fig. 1.4 but only a small fraction of this is intercepted by the Earth. Half the total energy emitted by the Sun is in the visible light range, with wavelengths from 0.4 to 0.7 /m. The rest arrives as ultraviolet or infrared waves, from 0.25 up to 3.0 /m.

The maximum rate of the Sun's emission (10 500kWm-2) occurs at 0.5 /m wavelength in the visible light range. Although there are changes in the solar output

Table 1.4

Pressure (mb) 1005 850 750

Specific humidity q (gkg-1) 14.2 12.4 9.5

SHAp 2061.5 1095.0

700

620

600

500

400

250

7.0

6.3

5.6

3.8

1.7

0.2

50

80

20

100

100

150

8.25

6.65

5.95

4.70

2.75

0.95

Fig. 1.4 Solar radiation.

associated with the occurrence of sunspots and solar flares, these are disregarded in assessing the amount of energy received by the Earth. The total solar radiation received in unit time on unit area of a surface placed at right angles to the Sun's rays at the Earth's mean distance from the Sun is known as the solar constant. The average value of the solar constant is 1.39kWm-2 (1.99calcm-2 min-2).

(b) Distance from the Sun. The distance of any point on the Earth's surface from the Sun is changing continuously owing to the Earth's eccentric orbit. The Earth is nearest the Sun in January at perihelion and furthest from it in July at aphelion. The solar constant varies accordingly.

(c) Altitude of the Sun. The Sun's altitude above the horizon has a marked influence on the rate of solar radiation received at any point on the Earth. The factors determining the Sun's altitude are latitude, season and time of day.

(d) Length of day. The total amount of radiation falling on a point of the Earth's surface is governed by the length of the day, which itself depends on latitude and season.

1.3.2.1 Atmospheric effects on solar radiation

The atmosphere has a marked effect on the energy balance at the surface of the Earth. In one respect it acts as a shield protecting the Earth from extreme external influences, but it also prevents immediate direct loss of heat. Thus it operates as an energy filter in both directions. The interchanges of heat between the incoming solar radiation and the Earth's surface are many and complex. There is a loss of energy from the solar radiation as it passes through the atmosphere known as attenuation. Attenuation is brought about in three principal ways as follows.

(a) Scattering. About 9 per cent of incoming radiation is scattered back into space through collisions with molecules of air or water vapour. A further 16 per cent are also scattered, but reach the Earth as diffuse radiation, especially in the shorter wavelengths, giving the sky a blue appearance.

(b) Absorption. Fifteen per cent of solar radiation is absorbed by the gases of the atmosphere, particularly by the ozone, water vapour and carbon dioxide. These gases absorb wavelengths of less than 0.3 /m only, and so very little of this radiation penetrates below an altitude of 40 km.

(c) Reflection. On average, 33 per cent of solar radiation is reflected from clouds and the ground back into space. The amount depends on the albedo (a) of the reflecting surfaces. White clouds and fresh white snow reflect about 90 per cent of the radiation (r = 0.9), but a dark tropical ocean under a high sun absorbs nearly all of it (a ^ 0). Between these two extremes is a range of surface conditions depending on roughness, soil type and water content of the soil. The albedo of the water surface of a reservoir is usually assumed to be 0.05, and of a short grass surface, 0.25.

1.3.2.2 Net radiation

As a result of the various atmospheric losses, only about 43 per cent of solar (shortwave) radiation reaches the Earth's surface, where most is absorbed and heats the land and oceans. The Earth itself radiates energy in the long-wave range (Fig. 1.5) and this long-wave radiation is readily absorbed by the atmosphere. The Earth's surface emits more than twice as much energy in the infrared range as it receives in short-wave solar radiation.

The balance between incoming and outgoing radiation varies from the Poles to the Equator. There is a net heat gain in equatorial regions and a net heat loss in polar regions. Hence, heat energy travels through circulation of the atmosphere from lower to higher latitudes. Further variations occur because the distribution of the continents and oceans leads to differential heating of land and water.

Fig. 1.5 Solar and terrestrial radiation.

Table 1.5 Average radiation values for selected latitudes (Wm 2)

July season

January season

Equator 30° S

250 280 170

210 240 220

190 240 230

-120 70 90

The amount of energy available at any particular point on the Earth's surface for heating the ground and lower air layers, and for the evaporation of water, is called the net radiation.

The net radiation RN may be defined by the equation: Rn = Ra - a(Ra) + Rl - Ro where Ra is the incoming short-wave (solar) radiation, a is the albedo, Ri is the incoming long-wave radiation and Ro is the outgoing long-wave radiation.

Incoming long-wave radiation comes from clouds (from absorbed solar radiation), and this has the following effects in the net radiation equations. In clear conditions, Rl ^ (0.6 to 0.8)Ro, thus Ri — Ro gives a net loss of long-wave radiation. For cloudy conditions, Ri ^ Ro and Ri — Ro becomes 0.

More significant are diurnal variations in net radiation, which is the primary energy source for evaporation. At night, S = 0 and Ri is smaller or negligible so that RN ^ Ro. In other words, net radiation is negative, and there is a marked heat loss, which is particularly noticeable when the sky is clear.

Some average values of solar (Ra), terrestrial (Ro) and net (RN) radiation for points on the earth's surface are given in Table 1.5.

1.4 Evaporation

Evaporation is the primary process of water transfer in the hydrological cycle. The oceans contain 95 per cent of the Earth's water and constitute a vast reservoir that remains comparatively undisturbed. From the surface of the seas and oceans, water is evaporated and transferred to temporary storage in the atmosphere, the first stage in the hydrological cycle.

1.4.1 Factors affecting evaporation

To convert liquid water into gaseous water vapour at the same temperature a supply of energy is required (in addition to that possibly needed to raise the liquid water to that temperature). The latent heat of vaporization (2.6 x 106Jkg-1) must be added to the liquid molecules to bring about the change of state. The energy available for evaporation is the net radiation obtaining at the water surface and is governed by local conditions of solar and terrestrial radiation.

The rate of evaporation is dependent on the temperature at the evaporating surface and that of the ambient air. It also depends on the vapour pressure of the existing water vapour in the air, since this determines the amount of additional water vapour that the air can absorb. From the saturation vapour pressure and air temperature relationship shown in Fig. 1.3, it is clear that the rate of evaporation is dependent on the saturation deficit. If the water surface temperature, Ts, is equal to the air temperature, Ta, then the saturated vapour pressure at the surface, es, is equal to ea. The saturation deficit of the air is given by (es - ed), where ed is the measure of the actual vapour pressure of the air at Ta.

As evaporation proceeds, the air above the water gradually becomes saturated and, when it is unable to take up any more moisture, evaporation ceases. The replacement of saturated air by drier air would enable evaporation to continue. Thus, wind speed is an important factor in controlling the rate of evaporation. The roughness of the evaporating surface is a subsidiary factor in controlling the evaporation rate because it affects the turbulence of the air flow.

In summary, evaporation from an open water surface is a function of available energy, the net radiation, the temperatures of surface and air, the saturation deficit and the wind speed. The evaporation from a vegetated surface is a function of the same meteorological variables, but it is also dependent on the presence of negative pressure potential (Section 6.2) within the soil or regolith. From a land surface, it is a combination of the evaporation of liquid water from precipitation collected on the land surface, from wetted vegetation surfaces and the transpiration of water by plants. Methods for the measurement of evaporation quantities are presented in detail in Chapter 4, and methods of analysis in Chapter 10.

1.5 Precipitation

The moisture in the atmosphere, although forming one of the smallest storages of the Earth's water, is the most vital source of fresh water for mankind. Water is present in the air in its gaseous, liquid and solid states as water vapour, cloud droplets and ice crystals, respectively.

The formation of precipitation from the water as it exists in the air is a complex and delicately balanced process. If the air was pure, condensation of the water vapour to form liquid water droplets would occur only when the air became greatly supersaturated. However, the presence of small airborne particles called aerosols provides nuclei around which water vapour in normal saturated air can condense. Many experiments, both in the laboratory and in the open air, have been carried out to investigate the requisite conditions for the change of state. Aitken (Mason, 1975) distinguished two main types of condensation nuclei: hygroscopic particles having an affinity for water vapour, on which condensation begins before the air becomes saturated (mainly salt particles from the oceans); and non-hygroscopic particles needing some degree of supersaturation, depending on their size, before attracting condensation. This latter group derives from natural dust and grit from land surfaces and from man-made smoke, soot and ash particles.

Condensation nuclei range in size from a radius 10-3 /xm for small ions to 10 /xm for large salt particles. The concentration of aerosols in time and space varies considerably. A typical number for the smallest particles is 40 000 per cm3, whereas for giant nuclei of more than 1 /m radius there might be only 1 per cm3. Large hygroscopic salt nuclei are normally confined to maritime regions, but the tiny particles called Aitken nuclei can travel across continents and even circumnavigate the Earth. Although condensation nuclei are essential for widespread condensation of water vapour, only a small fraction of the nuclei present in the air take part in cloud droplet formation at any one time.

Other conditions must be fulfilled before precipitation occurs. First, moist air must be cooled to near its dew point. This can be brought about in several ways as follows.

(a) By an adiabatic expansion of rising air. A volume of air may be forced to rise by an impeding mountain range. The reduction in pressure causes a lowering of temperature without any transference of heat.

(b) By a meeting of two very different air masses. For example, when a warm, moist mass of air converges with a cold mass of air, the warm air is forced to rise and may cool to the dew point. Any mixing of the contrasting masses of air would also lower the overall temperature.

(c) By contact between a moist air mass and a cold object such as the ground.

Once cloud droplets are formed, their growth depends on hygroscopic and surface tension forces, the humidity of the air, rates of transfer of vapour to the water droplets and the latent heat of condensation released. A large population of droplets competes for the available water vapour and so their growth rate depends on their origins and on the cooling rate of air providing the supply of moisture (Fig. 1.6).

Fig. 1.6 Comparative sizes, concentrations and terminal falling velocities of some particles involved in condensation and precipitation processes, where r = radius (^m); n = number per dm3 (I03 cm3); V = terminal velocity (cm s_'). (Reproduced from B.J. Mason (1975) Clouds, Rain and Rainmaking, 2nd edn, by permission of Cambridge University Press.)

Fig. 1.6 Comparative sizes, concentrations and terminal falling velocities of some particles involved in condensation and precipitation processes, where r = radius (^m); n = number per dm3 (I03 cm3); V = terminal velocity (cm s_'). (Reproduced from B.J. Mason (1975) Clouds, Rain and Rainmaking, 2nd edn, by permission of Cambridge University Press.)

Fig. 1.7 Frontal weather conditions, showing cloud and precipitation around cold and warm fronts. (Reproduced with permission from I. Strangeways (2007) Precipitation: Theory, Measurement and Distribution, Cambridge University Press, Cambridge.)

The mechanism becomes complicated when the temperature reaches freezing point. Pure water can be supercooled to about -40° C (233 K) before freezing spontaneously. Cloud droplets are unlikely to freeze in normal air conditions until cooled below -10°C (263K) and commonly exist down to -20°C (253K). They freeze only in the presence of small particles called ice nuclei, retaining their spherical shape and becoming solid ice crystals. Water vapour may then be deposited directly on to the ice surfaces. The crystals grow into various shapes depending on temperature and the degree of supersaturation of the air with respect to the ice.

Condensed water vapour appears in the atmosphere as clouds in various characteristic forms; a standard classification of clouds is shown in Fig. 1.7. The high clouds are composed of ice crystals, the middle clouds of either water droplets or ice crystals, and the low clouds mainly of water droplets, many of them supercooled. Clouds with vigorous upwards vertical development, such as cumulonimbus, consist of cloud droplets in their lower layers and ice crystals at the top.

1.5.1 Theories of raindrop growth

Considerable research has been carried out by cloud physicists on the various stages involved in the transference of atmospheric water vapour into precipitable raindrops or snowflakes. A cloud droplet is not able to grow to raindrop size by the simple addition of water vapour condensing from the air. It is worth bearing in mind that one million droplets of radius 10 fim are equivalent to a single small raindrop of radius 1 mm. Fig. 1.6 shows the principal characteristics of nuclei, cloud droplets and raindrops.

Cloud droplets can grow naturally to about 100 fim in radius, and although tiny drops from 100 to 500 fim may, under very calm conditions, reach the ground, other factors are at work in forming raindrops large enough to fall to the ground in appreciable quantities. There are several theories of how cloud droplets grow to become raindrops, and investigations into the details of several proposed methods continue to claim the attention of research workers.

The Bergeron process, named after the famous Norwegian meteorologist, requires the coexistence in a cloud of supercooled droplets and ice particles and a temperature less than 0° C (273 K). The air is saturated with respect to water but supersaturated with respect to ice. Hence water vapour is deposited on the ice particles to form ice crystals.

The air then becomes unsaturated with respect to water so droplets evaporate. This process continues until either all the droplets have evaporated or the ice crystals have become large enough to drop out of the cloud to melt and fall as rain as they reach lower levels. Thus the crystals grow at the expense of the droplets. This mechanism operates best in clouds with temperatures in the range -10 to -30° C (263-243 K) with a small liquid water content.

Growth by collision: in clouds where the temperature is above 0° C (273 K), there are no ice particles present and cloud droplets collide with each other and grow by coalescence. The sizes of these droplets vary enormously and depend on the size of the initial condensation nuclei. Larger droplets fall with greater speeds through the smaller droplets with which they collide and coalesce. As larger droplets are more often formed from large sea-salt nuclei, growth by coalescence operates more frequently in maritime than in continental clouds. In addition, as a result of the dual requirements of a relatively high temperature and generous liquid water content, the growth of raindrops by coalescence operates largely in summer months in low-level clouds.

When cloud temperatures are below 0° C (273 K) and the cloud is composed of ice particles, their collision causes growth by aggregation to form snowflakes. The most favourable clouds are those in the 0 to -4° C (269 K) range and the size of snowflakes decreases with the cloud temperature and water content.

Growth by accretion occurs in clouds containing a mixture of droplets and ice particles. Snow grains, ice pellets or hail are formed as cloud droplets fuse on to ice particles. Accretion takes place most readily in the same type of cloud that favours the Bergeron process, except that a large content of liquid water is necessary for the water droplets to collide with the ice particles.

Even when raindrops and snowflakes have grown large enough for their gravity weight to overcome up-draughts of air and fall steadily towards the ground, their progress is impeded by changing air conditions below the clouds. The temperature may rise considerably near the Earth's surface and the air may become unsaturated. As a result snowflakes usually melt to raindrops and the raindrops may evaporate in the drier air. On a summer's day it is not uncommon to see cumulus clouds trailing streams of rain which disappear before they reach the ground. With dry air below a high cloud base of about 3 km, all precipitation will evaporate. Hence it is rare to see rainfall from altocumulus, altostratus and higher clouds (see Fig. 1.7). Snowflakes rarely reach the ground if the surface air temperature is above 4° C, but showers of fine snow can occur with the temperature as high as 7° C, if the air is very dry.

Further explanation of the processes involved in raindrop formation is given in Sumner (1988) and Strangeways (2007).

1.6 Weather patterns producing precipitation

The main concern of the meteorologist is an understanding of the general circulation of the atmosphere with the aim of forecasting the movements of pressure patterns and their associated winds and weather. It is sufficient for the hydrologist to be able to identify the situations that provide the precipitation, and for the practising civil engineer to keep a 'weather eye' open for adverse conditions that may affect his site work.

The average distribution and seasonal changes of areas of high atmospheric pressure (anticyclones) and of low-pressure areas (depressions) can be found in most good atlases. Associated with the location of anticyclones is the development of homogeneous air masses. A homogeneous air mass is a large volume of air, generally covering an area greater than 1000 km in diameter, which shows little horizontal variation in temperature or humidity. It develops in the stagnant conditions of a high-pressure area and takes on the properties of its location (known as a source region). In general, homogeneous air masses are either cold and stable, taking on the characteristics of the polar regions from where they originate, or they are warm and unstable, revealing their tropical source of origin. Their humidity depends on whether they are centred over a large continent or over the ocean. The principal air masses are summarized in Table 1.6. Differences in atmospheric pressure cause air masses to move from high- to low-pressure regions and they become modified by the environments over which they pass. Although they remain homogeneous, they may travel so far and become so modified that they warrant reclassification. For example, when polar maritime air reaches the British Isles from a south-westerly direction, having circled well to the south over warm subtropical seas, its character will have changed dramatically.

Precipitation can come directly from a maritime air mass that cools when obliged to rise over mountains in its path. Such precipitation is known as orographic rainfall (or snowfall, if the temperature is sufficiently low), and is an important feature of the western mountains of the British Isles, which lie across the track of the prevailing winds bringing moisture from the Atlantic Ocean. Orographic rain falls similarly on most hills and mountains in the world, with similar locational characteristics, though it may occur only in particular seasons.

When air is cooled as a result of the converging of two contrasting air masses, it can produce more widespread rainfall independent of surface land features. The boundary between two air masses is called a frontal zone. It intersects the ground at the front, a band of about 200 km across. The character of the front depends on the difference between the air masses. A steep temperature gradient results in a strong or active front and much rain, but a small temperature difference produces only a weak front with less or even no rain. The juxtaposition of air masses across a frontal zone gives rise to two principal types of front according to the direction of movement.

Fig. 1.8 illustrates cloud patterns and the weather associated firstly with a warm front, in which warm air is replacing cold air, and secondly with a cold front, in which cold air is pushing under a warm air mass. In both cases, the warm air is made to rise and hence cool, and the condensation of water vapour forms characteristic clouds

Table 1.6 Classification of air masses

Air mass

Source region

Properties of source

Polar maritime (Pm) Polar continental (Pc)

Arctic or Antarctic (A)

Tropical maritime (Tm)

Tropical continental (Tc)

Oceans; 50° latitude Continents in vicinity of Arctic

Circle; Antarctica Arctic Basin and Central Antarctica in winter Sub-tropical oceans

Deserts in low latitude; primarily the Sahara and Australian deserts

Cool, rather moist, unstable Cool, dry, stable

Very cold, dry, stable

Warm and moist; unstable inversion common feature Hot and dry

Fig. 1.8 Frontal weather conditions, showing cloud and precipitation development around an occluded front. (Reproduced from Strangeways, I. (2007) Precipitation: Theory, Measurement and Distribution, Cambridge University Press, Cambridge.)

and rainfall. The precipitation at a warm front is usually prolonged with gradually increasing intensity. At a cold front, however, it is heavy and short-lived. Naturally, these are average conditions; sometimes no rain is produced at all.

Over the world as a whole there are distinctive regions between areas of high pressure where differing air masses confront each other. These are principally in the mid-latitudes between 30° and 60° in both hemispheres, where the main boundary, the polar front, separates air masses having their origins in polar regions from the tropical air masses.

In addition, there is a varying boundary between air masses originating in the northern and southern hemispheres known as the intertropical convergence zone (ITCZ). The seasonal migration of the ITCZ plays a large part in the formation of the monsoon rains in south-east Asia and in the islands of Indonesia.

Four major weather patterns producing precipitation have been selected for more detailed explanation.

1.6.1 Mid-latitute cyclones or depressions

Depressions are the major weather pattern for producing precipitation in the temperate regions. More than 60 per cent of the annual rainfall in the British Isles comes from such disturbances and their associated features. They develop along the zone of the polar front between the polar and tropical air masses. Knowledge of the growth of depressions, the recognition of air masses and the definition of fronts all owe much to the work of the Norwegian meteorologists Wilhelm and Jacob Bjerknes in the 1920s.

The main features in the development and life of a mid-latitude cyclone are shown in Fig. 1.9. The first diagram illustrates in plan view the isobars of a steady-state condition at the polar front between contrasting air masses. The succeeding diagrams show the sequential stages in the average life of a depression. A slight perturbation caused by irregular surface conditions, or perhaps a disturbance in the lower stratosphere, results in a shallow wave developing in the frontal zone. The initial wave, moving along the line of the front at 15-20 ms-1 (30-40 knots), may travel up to 1000 km without further development. If the wavelength is more than 500 km, the wave usually increases in amplitude, warm air pushes into the cold air mass and active fronts are formed.

Fig. 1.9 Life cycle of a model occluding depression. (Adapted from Met Office (1962) A Course in Elementary Meteorology, Her Majesty's Stationary Office.)

As a result, the air pressure falls and a 'cell' of low pressure becomes trapped within the cold air mass. Gradually the cold front overtakes the warm front, the warm air is forced aloft, and the depression becomes occluded. The low-pressure centre then begins to fill and the depression dies as the pressure rises. On average, the sequence of growth from the first perturbation of the frontal zone to the occlusion takes 3-4 days. Precipitation usually occurs along the fronts and, in a very active depression, large amounts can be produced by the occlusion, especially if its speed of passage is retarded by increased friction at the Earth's surface. At all stages, orographic influences can increase the rainfall as the depression crosses land areas. A range of mountains can delay the passage of a front and cause longer periods of rainfall. In addition, if mountains delay the passage of a warm front, the occlusion of the depression may be speeded up.

1.6.2 Waves in the easterlies and tropical cyclones

Small disturbances are generated in the trade wind belts in latitudes 5-25° both north and south of the Equator. Irregular wind patterns showing as isobaric waves on a weather map develop in the tropical maritime air masses on the equatorial side of the subtropical high-pressure areas. They have been studied most in the Atlantic Ocean to the north of the South American continent. A typical easterly wave is shown in Fig. 1.10. A trough of low pressure is shown moving westwards on the southern flanks of the Azores anticyclone. The length of the wave extends over 15-20° longitude (1500km) and, moving with an average speed of 6.7ms-1 (13knots), takes 3-4 days to pass. The weather sequence associated with the wave is indicated beneath

Fig. 1.10 A wave in the easterlies. Weather sequence: 1 - small Cu, no pp.; 2 - Cu, a few buildups, haze, no pp.; 3 - larger Cu, Ci and Ac, better visibility, pr ... pr ... pr (showers); 4 - very large Cu, overcast Ci Ac, prpr or rr (continuous rain); 5 - Cu and Cb, Sc, As, Ac, Ci, pRpR (heavy showers), (thunderstorm); 6 - large Cu, occasional Cb, some Sc, Ac, Ci, pRpr - prpr; Cu = Cumulus; Cb = Cumulonimbus; Ci = Cirrus; Ac = Altocumulus; As = Altostratus; Sc = Stratucumulus.

Fig. 1.10 A wave in the easterlies. Weather sequence: 1 - small Cu, no pp.; 2 - Cu, a few buildups, haze, no pp.; 3 - larger Cu, Ci and Ac, better visibility, pr ... pr ... pr (showers); 4 - very large Cu, overcast Ci Ac, prpr or rr (continuous rain); 5 - Cu and Cb, Sc, As, Ac, Ci, pRpR (heavy showers), (thunderstorm); 6 - large Cu, occasional Cb, some Sc, Ac, Ci, pRpr - prpr; Cu = Cumulus; Cb = Cumulonimbus; Ci = Cirrus; Ac = Altocumulus; As = Altostratus; Sc = Stratucumulus.

the diagram. In the tropics, the cloud-forming activity from such disturbances is vigorous and subsequent rainfall can be very heavy: up to 300 mm may fall in 24 h.

As in mid-latitudes, the wave may simply pass by and gradually die away, but the low pressure may deepen with the formation of a closed circulation with encircling winds. The cyclonic circulation may simply continue as a shallow depression giving increased precipitation but nothing much else. However, rapidly deepening pressure below 1000 mb usually generates hurricane-force winds blowing round a small centre of 30-50 km radius, known as the eye. At its mature stage, a hurricane centre may have a pressure of less than 950 mb. Eventually the circulation spreads to a radius of about 300 km and the winds decline. Copious rainfall can occur with the passage of a hurricane; record amounts have been measured in the region of Southeast Asia, where the effects of the storms have been accentuated by orography. However, the rainfall is difficult to measure in such high winds. In fact, slower moving storms usually give the higher records. Hurricanes in the region of Central America often turn northwards over the United States and die out over land as they lose their moisture. On rare occasions disturbances moving along the eastern coastal areas of the United States are carried into westerly air-streams and become vigorous mid-latitude depressions.

Hurricanes tend to be seasonal events occurring in late summer when the sea temperatures in the areas where they form are at a maximum. They are called typhoons in the China Seas and cyclones in the Indian Ocean and off the coasts of Australasia. These tropical disturbances develop in well-defined areas and usually follow regular tracks; an important fact when assessing extreme rainfalls in tropical regions (McGregor and Nieuwolt, 1998).

1.6.3 Convectional precipitation

A great deal of the precipitation in the tropics is caused by local conditions that cannot be plotted on the world's weather maps. When a tropical maritime air mass moves over land at a higher temperature, the air is heated and forced to rise by convection. Very deep cumulus clouds form, becoming cumulonimbus extending up to the tropopause. Fig. 1.11 shows the stages in the life cycle of a typical cumulonimbus. Sometimes these occur in isolation, but more usually several such convective cells grow together and the sky is completely overcast.

The development of convective cells is a regular daily feature of the weather throughout the year in many parts of the tropics, although they do not always provide rain. Cumulus clouds may be produced but evaporate again when the air ceases to rise. With greater vertical air velocities, a large supply of moisture is carried upwards. As it cools to condensation temperatures, rainfall of great intensity occurs. In extreme conditions, hail is formed by the sequential movement of particles up and down in the cloud, freezing in the upper layers and increasing in size by gathering up further moisture. As the rain and hail fall, they cause vigorous down draughts, and when these exceed the vertical movements, the supply of moisture is reduced, condensation diminishes and precipitation gradually dies away. Thunder and lightning are common features of convectional storms with the interaction of opposing electrical charges in the clouds. The atmospheric pressure typically is irregular during the course of a storm.

Convectional activity is not confined to the tropics; it is a common local rain-forming phenomenon in higher latitudes, particularly in the summer. Recent studies have shown km 1

km 1

Mmmw*

(b) Mature

(a) Developing

(c) Dissipating

Fig. I.ll Convective cells - stages in the life cycle. Time scales: (a) approximately 20 min; (b) approx imately 20 min, heavy rain and hail, thunder may develop; (c) 30 min to 2 h, rainfall intensity decreasing. Total life cycle 1-2 h.^, ice; *, snow; • •• •, rain and hail; winds.

that convection takes place along frontal zones thus adding to rainfall intensities. Wherever strong convectional forces act on warm moist air, rain is likely to form and it is usually of high intensity over a limited area.

1.6.4 Monsoons

Monsoons are weather patterns of a seasonal nature caused by widespread changes in atmospheric pressure. The most familiar example is the monsoon of Southeast Asia where the dry, cool or cold winter winds blowing outwards from the Eurasian anticyclone are replaced in summer by warm or hot winds carrying moist air from the surrounding oceans being drawn into a low-pressure area over northern India. The seasonal movements of the ITCZ play a large part in the development and characteristics of the weather conditions in the monsoon areas. The circulation of the whole atmosphere has a direct bearing on the migration of the ITCZ, but in general the regularity of the onset of the rainy seasons is a marked feature of the monsoon. Precipitation, governed by the changing seasonal winds, can be caused by confrontation of differing air masses, low-pressure disturbances, convection and orographic effects. A map of the monsoon areas is shown in Fig. 1.12. Actual quantities of rain vary, but as in most tropical and semi-tropical countries, intensities are high (McGregor and Nieuwolt,

Further explanation of the processes producing the weather systems described is given in Holton (2004).

1.7 Climate

Following the appreciation of the meteorological mechanisms that affect evaporation and produce precipitation, it is pertinent to consider these hydrological processes

Fig. 1.12 Monsoon lands - pressure systems and winds.

on a longer time scale. Evaporation was presented as an instantaneous process. The precipitation-forming mechanisms extended into weather patterns that may last up to about a week. The study of climate is based on average weather conditions, specified usually by measures of temperature and precipitation over one or more months though other phenomena may also be aggregated. Statistics gathered for each month over a period of years and averaged give a representation of the climate of the location.

Table 1.7 Köppen climate classification

Estimated percentages

Land surfaces Total surface

Table 1.7 Köppen climate classification

Estimated percentages

Land surfaces Total surface

A

Tropical rain climates - forests

20

36

B

Arid climates

26

11

C

Warm temperate rain climates - trees

15

27

D

Boreal forest and snow climates

21

7

E

Treeless cold snow climates

17

19

The most renowned classification of climates is that of Koppen who categorized climates according to their effect on vegetation.

The major groupings are given in Table 1.7. Subdivisions of these main groups are defined by thresholds of temperature and rainfall values; the details are given in most books of climatology. Their geographical distribution is shown in Fig. 1.13, though Peel et al. (2007) provide a revised (colour) version of this map.

These broad definitions of climatic regions are built up from the instrumental records of observing stations which are thus providing sample statistics representing conditions over varying areas. Such meteorological records have only been made with any reliability since the advancing development of instruments in the seventeenth century (Manley, 1970), and world coverage was limited until the late nineteenth century.

Before instrumental records, knowledge of the climate of different regions has been built up by the study of what is now called proxy data. For example, in the UK the proportion of certain tree pollens found in layers of lake sediments or upland peats give indications of the existence of tree cover in earlier times. Similarly, varying layers of clays and silts in surface deposits, as in Sweden, help to differentiate between warm and cold periods. In the western United States, the study of growth rings in the trunks of very old trees, allow climatologists to extend climatic information to periods before instrumental observations. On the global scene, the analyses of deep-sea sediments and ice cores are of increasing importance in the assemblage of climate knowledge.

In addition, archaeological and historical records of transient events, such as the extent of sea ice round the Poles, the fluctuation of mountain glaciers and even the variation of man's activities in the extent of vine growing and the abundance of the wheat harvests, all contribute clues to the climate of former times.

The assimilation and interpretation of such variable information gathered worldwide has occupied climatologists for many years and a broadly agreed sequence of climatic events has been established, aided by the findings of the geologists. However, the worldwide coverage of climatic information before this century was far from representative of all land regions and even less was known of the much larger oceanic areas. The recent concern over man-made changes in the composition of the atmosphere and the increasing ability to model changes in climate has led scientists to study the dynamic components contributing to climate as distinct from current weather, and the resultant impacts on hydrology (Bates et al., 2008). Detailed discussion of climate change impacts on hydrology are presented within Chapter 19.

A-C boundaries E boundaries

BW boundaries internal boundaries between A climates

B S boundaries ------ internal boundaries between C climates

C-D boundaries ----internal boundaries between D climates

Fig. 1.13 Köppen's world classification of climates.

Note

1 Runoff is the river discharge (m3 s-1) per unit catchment area, hence has the units of m s-1 or mmh-1.

References

Bates, B. C., Kundzewicz, Z. W., Wu, S. and Palutikof, J. P. (eds) (2008) Climate Change and Water. Technical Paper of the Intergovernmental Panel on Climate Change, IPCC Secretariat, Geneva, 210 pp.

Beven, K. J. (2004) Robert Horton's perceptual model of infiltration. Hydrological Processes 18, 3447-3460.

Beven, K. J. (2006) Benchmark Papers in Storm Runoff Generation. IAHS Press, Wallingford.

Biswas, A. K. (1970) History of Hydrology. North-Holland Pub. Co., Amsterdam, 336 pp.

Brustseart, W. (2005) Hydrology - An Introduction. Cambridge University Press, Cambridge.

Bunting, B. T. (1961) The role of seepage moisture in soil formation, slope development and stream initiation. American Journal of Science 259, 503-518.

Dooge, J. C. I. (1959) Un bilan hydrologique au XVIIe siecle. Houille Blanche, 14e annee No. 6, 799-807.

Dunne, T. (1978) Field studies of hillslope flow processes. In: Kirkby, M. J. (ed.) Hillslope Hydrology, John Wiley & Sons, pp. 227-294.

Gleick, P. H. (1996) Water resources. In: Schneider, S. H. (ed.) Encyclopedia of Climate and Weather, Vol. 2. Oxford University Press, New York, pp. 817-823.

Holton, J. R. (2004) An Introduction to Dynamical Meteorology, 4th edn. Academic Press, San Diego, 535 pp.

Manley, G. (1970) The climate of the British Isles. In: Wallen, C. C. (ed.) Climates of Northern and Western Europe, Elsevier, Amsterdam, 81-133. (World Survey of Climatology, Vol. 5)

Mason, B. J. (1975) Clouds, Rain and Rainmaking, 2nd edn. Cambridge University Press, Cambridge, 189 pp.

McGregor, G. R. and Nieuwolt, S. (1998) Tropical Climatology: An Introduction to the Climates of the Low Latitudes. Wiley, Chichester, 339 pp.

Met Office (1978) A Course in Elementary Meteorology, 2nd edn. HMSO, London, 208 pp.

Peel, M. C., Finlayson, B. L. and McMahon, T. A. (2007). Updated world map of the KoppenGeiger climate classification. Hydrology and Earth System Science 1, 1633-1644.

Sklash, M. G. and Farvolden, R. N. (1979) The role of groundwater in storm runoff. Journal of Hydrology 43, 45-65.

Sklash, M. G., Beven, K. J., Gilman, K. and Darling, W. G. (1996) Isotope studies of pipeflow in Plynlimon, Wales, UK. Hydrological Processes 10, 921-944.

Strangeways, I. (2007) Precipitation: Theory, Measurement and Distribution. Cambridge University Press, Cambridge.

Sumner, G. (1988) Precipitation: Process and Analysis. John Wiley & Sons, Chichester, 455 pp.

Weiler, M. and McDonnell, J. (2004) Virtual experiments: a new approach for improving process conceptualization in hillslope hydrology. Journal of Hydrology 285, 3-18.

Weyman, D. R. (1970) Throughflow on hillslopes and its relation to the stream hydrograph. International Association of Scientific Hydrology Bulletin 15, 25-33.

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