hydrological cycle

hydrological cycle The hydrological cycle describes the various routes by which water moves from the atmosphere and oceans to the land surface and returns to the oceans via surface and subsurface pathways (runoff). It is the central concept in hydrology, and hydrologists strive both to understand and to predict the fate of water on the land at a variety of spatial scales.

At the global scale, 97.2 per cent of water resides in the oceans. Of the remaining 2.8 per cent, which is made up of both fresh and brackish water, 77 per cent is locked up in ice sheets and glaciers, 22 per cent is stored beneath the Earth's surface as groundwater, and the remaining 1 per cent is stored in the atmosphere, rivers, lakes, inland seas, and soil. The hydrological cycle is driven by the evaporation of water from both land and ocean surfaces and the lateral movement of moisture-laden air to the continents. Some 419 cubic kilometres (km³) of water are evaporated from the oceans every year, and about 69 km³ of water are evaporated from the continents. Precipitation on the continents amounts to about 106 km³ and run-off is about 37 km³. The oceans receive proportionately less rainfall (382 km³) than they supply to the atmosphere by evaporation, and the continents therefore receive proportionately more than they supply to the atmosphere by evaporation. The inputs to and outputs from the continents are in balance. The following expression shows this balance as a simple equation (quantities are in km³):

Evaporation	+	Run-off	=	Precipitation.
69	+	37	=	106

This is referred to by hydrologists as the water balance. From this expression we can calculate the run-off ratio (the ratio of run-off to precipitation) which, at the global scale, is 35 per cent. As we shall see later, the water balance varies significantly at the surface of the Earth because of regional differences in precipitation and evaporation.

At the regional scale, there are significant differences in the amount of precipitation received by the continents. This results from variations in the global movement of moisture-laden air masses. Two regions receive more than the average annual global precipitation of about 85–90 cm. These are the low-latitude equatorial regions (e.g. Amazonia and Indonesia) and the west coasts of middle-latitude continents (e.g. Western Europe, Alaska, and Western Canada). There are two regions which receive less than the annual average global precipitation. These include areas where semi-permanent high-pressure systems develop (e.g. the arid and semi-arid regions of the south-west United States and the Saharan region of North Africa) and the continental interiors of high-latitude regions (e.g. central Siberia, central Canada); the low-temperature air masses in these regions have a very low moisture content.

The global inequalities in the distribution of precipitation and in evaporation have a profound effect on the water balance at the regional scale. Table 1 gives annual precipitation, evaporation, and run-off data for three geographical regions of Eastern Europe. These data show that there is an unequal distribution of precipitation at the regional scale but, more significantly, also show us that the run-off ratio varies with latitude. In general, high run-off ratios (over 50 per cent) are experienced in cool high-latitude regions whereas low run-off ratios are found in the relatively warmer low-latitude regions.

At the scale of the catchment or drainage-basin, the hydrological cycle comprises one major input, six stores, seven transfer processes, and three outputs (Fig. 1).

Catchment inputs; precipitation. Precipitation occurs in a variety of forms, including fog, drizzle, rain, sleet, hail, and snow. The most important input in high latitudes and at high altitudes is snow. Snow will often accumulate and can eventually compact to form glaciers and ice caps. In this case, the input is released slowly into the hydrological cycle and the output via river run-off is controlled by temperature. The melting of ice and snow reach a seasonal maximum in the summer months, and river discharge is high at this time of the year. At the daily timescale, the highest run-off is recorded in the late afternoon following a rise in temperature. The most common form of precipitation outside high latitudes and high altitudes is rain. Hydrologists are concerned with the amount of rain received by a drainage basin (usually expressed as rainfall depth in millimetres or centimetres) and the intensity at which it falls, measured in millimetres per hour (mm hr⁻¹). Rainfall volume and intensity are important in controlling the output from the catchment. High-intensity storms are usually associated with convectional instability, which causes thunderstorms. These are often local in nature but can lead to excessive amounts of precipitation in local areas and can cause severe flooding. Many of these storms occur in hot arid and tropical environments. Since the soils of arid environments are often thin and compacted, flooding is exacerbated by the inability of the catchment soils to absorb the high input. In tropical areas, high-intensity rainfalls are commonly associated with landslides, mudflows, and other mass-movement hazards.

Table 1. Annual water balance of three geographical regions of Eastern Europe

	Precipitation	Evaporation	Run-off	Run-off ratio
(cm)	(cm)	(cm)	(%)
Tundra	45	11	34	76
Mixed forest	58	37	21	36
Semi-desert	20	19	1	5

Catchment stores. Water moving through the hydrological cycle can be stored in one or six locations (Fig. 1). Above the ground surface are the interception, vegetation, surface, and channel stores. Below the ground surface are the soil and groundwater stores. Vegetation acts both passively and actively in storing water. Water in the interception store is held on the canopy surface, which includes the leaves, stems, and branches of growing vegetation. In the vegetation is the water held at any one time within the vegetation. Vegetation is usually replenished directly from the soil-water store. The quantitative importance of the interception and vegetation stores depends on the type of vegetation. In general, forests have a higher storage capacity than crops and grasslands, although there are significant differences in the volume stored, depending, for example, on the age and species of the trees. The species is particularly important in relation to the seasonal variation in storage volume, since conifers retain a leaf cover throughout the year, whereas deciduous forests lose their leaves in the winter. Recent studies in Europe have shown that oak forests can store as much as 20 mm per month whereas Douglas Fir forests can store more than 40 mm per month. Seasonality in vegetation and interception storage is also found in cropped agricultural systems. The importance of canopy interception increases in areas where rainfall occurs as light frequent showers. In mixed deciduous forests, 60–70 per cent of incoming precipitation can be stored in the canopy for storms of 1–2 hours duration. This decreases to less than 20 per cent in storms of over 12 hours duration. Snow can also be stored in substantial quantities in the canopy, and its transfer to the ground is largely controlled by wind and exposure.

The surface store accounts for the volume of water which can be stored in micro-scale hollows in the ground surface, which, although they are most apparent in cultivated landscapes, are also known to occur naturally. Any regular or irregular depression in the ground surface can prevent water flowing downslope towards the river channel. These depressions include plough furrows running at right angles to the dominant slope direction or irregularities left after harrowing. The presence of shallow depressions in the ground surface allows time for water to percolate into the soil and reduces the volume and speed of flow across the slope. The importance of small depressions in allowing the soil-water store to be recharged has long been appreciated in arid and semi-arid environments, where contour ploughing and bunding (creation of embankments) are commonly used as a means of sustaining rain-fed agriculture by inducing soil-water recharge. Although we usually assume that surface storage is a small-scale phenomenon, it may also be of significance at the larger scale, where water may enter the soil and groundwater stores from deflation hollows (depressions left from the activity of wind in arid environments) and kettle holes (depressions left by the melting of stagnant ice during deglaciation).

The channel store measures the amount of water held within the river channel at any particular instant. Although recognized by hydrologists as one type of store, it is a relatively minor component in the hydrological cycle.

Beneath the ground surface are the two major storage components of soil and groundwater. Precipitation enters the ground surface and moves downwards through the soil towards the water table, which marks the upper limit of the zone of saturation (the groundwater zone). Water is held against the influence of gravity in the soil zone because of the attraction of water molecules for each other and for fine soil particles. There are three types of water held within a soil. In dry soils, a thin film of hygroscopic water is retained around individual soil particles. It is effectively unavailable to plants, which cannot overcome the surface tensions involved in order to release the water from the soil. Other water molecules link up with the hygroscopic water and are held less strongly in the soil because the water–water bond is less strong than the soil–water bond. This is known as capillary water. The third type of water is not held by the soil but moves slowly through it under the influence of gravity. This is gravitational water, which drains slowly downwards towards the groundwater zone. A soil that has its full quota of capillary water is said to be at field capacity (i.e. it contains the maximum amount of water that a soil can hold against the influence of gravity). The amount of water that can be stored in the subsurface zone depends upon the size and arrangement of particles. A soil composed predominantly of clay and fine silt particles, for example, will have a higher storage capacity (porosity) than a sandy soil. In the soil zone the actual amount of water stored will depend on the porosity of the soil and the amount lost by evaporation, drainage, and plant uptake. Soils therefore usually hold less than their maximum potential. In contrast, groundwater is permanently saturated and the porosity is controlled by the void spaces in the rock. These voids include the intergranular spaces in the intact rock as well as the fractures and joints that develop as a result of earth movement. Groundwater can occur at great depths in the subsurface zone and in near-surface horizons where impermeable strata prevent the downward movement of water. In this latter case, a perched water table will develop.

Catchment transfers. Water transfer through the hydrological cycle is controlled through seven mechanisms. Three of these, throughfall and stemflow, surface run-off, and channel flow occur above the ground surface; a fourth, infiltration, controls the transfer of water from the ground surface to the subsurface zone; and the remaining three, percolation, throughflow, and groundwater flow, occur beneath the ground surface (Fig. 1).

Throughfall and stemflow are terms that describe the movement of water between the interception store and the ground surface. Throughfall refers to the free fall of water to the ground after it has collected on vegetation surfaces. This water commonly has a much larger drop size than rain and is therefore more destructive when it hits the ground. Stem flow describes the flow of water along branches and down the trunks and stems of growing vegetation to reach the ground surface at the base of the plant. Stemflow and throughfall transfers therefore play an important role in modifying the spatial distribution and intensity characteristics of rainfall reaching the ground surface.

Surface run-off, sometimes called overland flow, is an important transfer mechanism in the hydrological cycle. As more water is transferred to river channels by this process, so the magnitude of a flood from the same volume of rainfall increases. This is partly because the velocity of water across the ground surface is much higher than through the soil. Surface run-off velocities of over 0.14 metres per second (m s⁻¹) have been observed in comparison with lateral water velocities through the soil of only 1 to 80 m s⁻¹ × 10⁻⁶. The amount of rainfall that becomes surface run-off is controlled by one of two mechanisms. Where the intensity of rainfall is higher than the rate at which water can enter the ground (the infiltration rate), surface run-off will occur. This is a common mechanism by which surface run-off is generated in arid and semi-arid areas, where rainfall intensities are often high and infiltration rates are low. In humid environments, the soil may become saturated and rainfall cannot enter the ground surface. The British geomorphologist M. J. Kirby has used the analogy of a bottle to distinguish between these processes. We either have a bottle with a very narrow neck that fills slowly or a bottle that is already full. In either case, the net result is the generation of surface run-off.

Infiltration refers to the rate at which water enters into the soil profile. The rate of infiltration is clearly of significance to the generation of surface run-off as described above, and also to the recharge of the soil and groundwater stores. Of greatest significance in controlling infiltration are the nature and properties of the soil and the presence or absence of a vegetation cover. Soils that have recently been ploughed or have a high-density vegetation cover generally have a high infiltration rate. Grassland that has been compacted by the trampling of grazing animals or bare soils that have been beaten by high-intensity rainfall generally have low infiltration rates. The size of soil particles is also important. Soils with even small quantities of clay can form impermeable crusts owing to high-intensity rainfall. In arid environments, the presence of biofilms (thin algal crusts) may reduce infiltration substantially.

Percolation refers to the vertical movement of water through the soil profile, whereas throughflow (or interflow) refers to the downslope transfer of water through the soil towards river channels. Percolation occurs through the soil matrix (the main mass of the soil) and through small cracks and spaces that are created by burrowing animals, tree roots, or the presence of certain clay minerals that expand and contract upon wetting and drying. If the upper soil profile contains gravitational water with the soil at or near saturation, water will move downwards under the influence of gravity at a rate that is controlled by the characteristics of the soil. Where capillary water is held, movement will usually be from a zone of high moisture content to a zone of low moisture content. Because this movement is controlled by the capillary tension forces that hold water in the soil, it is more precise to say that water moves from a zone of low capillary tension to a zone of high capillary tension. This movement can occur both upwards and downwards within the soil profile.

Throughflow occurs when there are significant changes in the density of different layers within the soil horizon. High-density impermeable layers are commonly found within soil horizons that allow the build-up of a saturated wedge of water above the impermeable layer. On a slope, water will move laterally under the influence of gravity. The rate of water movement is again controlled by the number and size of the void spaces within the soil. Although the term ‘throughflow’ is generally used to describe the lateral movement of water through the soil matrix, in some environments small pipes and tunnels develop at the junction between a permeable and impermeable layer in which water can move more rapidly than through the soil matrix. Pipes and tunnels, in which water velocities approach that of overland flow, are common in areas of peat and also in areas with fine loess (wind-blown) silts, where tunnels of up to 4 m in diameter may develop.

Groundwater flow, like throughflow, is the slow downslope movement of saturated water under the influence of gravity within the saturated zone. Rates of movement are variable but are usually very slow. Sandstones typically have water velocities of between 0.28 × 10⁻⁶ and 0.28 × 10⁻² m s⁻¹, whereas in closely jointed limestones velocities may be as high as 0.14 m s⁻¹. In areas where the water table is confined between impermeable strata and the groundwater recharge area is at a higher elevation some distance away, water may be forced to the ground surface under pressure. This is artesian water, which is often a commercially important source for water supply.

Channel flow is the most efficient means of removing water from a catchment. At high water discharge, velocities can exceed of 3 m s⁻¹. The velocity of flow is controlled by the gradient of the channel, the shape of the channel, and the roughness caused by the presence of stones and vegetation. In both urban and rural areas, artificial drainage frequently increases the length of open channels and pipes. While these serve an important function in improving soil drainage for cultivation and in preventing flooding in urban areas, they are also perceived as posing an increased risk of flooding in downstream areas.

Catchment outputs. There are three catchment outputs, evaporation, transpiration, and run-off (Fig. 1). Evaporation is an output from the interception, soil surface, and soil water stores. Transpiration is an output from the vegetation store, usually via the stomata on the leaves of living vegetation. The vegetation stores can tap water directly from the soil water store. To an extent, transpiration rates are regulated by the biosphere. Evaporation and transpiration rates are primarily controlled by a number of other factors, not least of which is the availability of a moisture supply. Supply limitation is of little importance in controlling evaporation from open water bodies, but is of great significance in relation to evaporation from bare soils, from the interception store, and from the vegetation store. The most significant factors controlling evaporation and transpiration losses are the amount of incident solar radiation (the energy which drives evaporation and transpiration) and the vapour pressure of the air relative to saturation. This controls the capacity of the air to absorb more moisture from the surrounding environment. In arid and semi-arid environments, evaporation dominates the output from the hydrological cycle. Run-off describes the loss of water from the catchment. It is measured by hydrologists at a wide range of timescales from hours to years. Annual run-off is controlled by the balance between the input of precipitation and the combined outputs of evaporation and transpiration giving rise to differences in the run-off ratio described above. At shorter timescales, run-off is controlled by differences in the magnitude of the storage components within the hydrological cycle and by the transfer processes between the catchment and river channel.

Of major significance in hydrological research is the direct and indirect impact of human activity on the various storages and transfer processes illustrated in Fig. 1. The process of urbanization, for example, reduces the capacity of the interception and vegetation stores, increases the efficiency of surface run-off and channel flow transfers, and consequently increases run-off at the hourly and annual timescales. Forest clearance and cultivation have similar effects on the natural hydrological cycle, but pose additional threats to the environment by making large quantities of sediment available for erosion.

Ian D. L. Foster

Bibliography

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