Types of Erosion and Its Assessment

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Soil is a vital resource for the production of renewable resources for the necessities of human life, such as food and fiber. Soils, however, essentially are nonrenewable resources [1].

According to Golubev [2], the area of cultivated land in the world is 14.3 million km2. In cultivated areas, drastic changes in vegetation have occurred and instead of dense natural vegetation cover, bare soil often is exposed for most of the year with sparse crop vegetation existing for a few months. These changes in vegetation cover are the main reason for the increase of soil erosion on cropland as compared to that on natural landscapes. Results of computations by Golubev [2] show that soil erosion in the world is 5.5 times more than during the preagricultural period. According to Brown [3], the world is currently losing 23 billion tonnes of soil from cropland in excess of new soil formation each year [4]; therefore, accelerated soil erosion is a serious problem to consider for the development of a sustainable agriculture. Other environmental problems caused by severe soil erosion are reservoir sedimentation, which results in a lowering of the available surface water resources, and nonpoint-source pollution due to sediment transport phenomena.

On a global scale, even if the mean annual sediment yield estimate is based on the available suspended sediment transport measurements, Walling and Webb [5] gave a reliable assessment of the global pattern of water erosion. This assessment established that the semiarid and semihumid areas of the world (China, India, western United States, and Mediterranean lands) are the most vulnerable to soil erosion.

Soil erosion losses are often due to a few severe storms with high rainfall intensity and/or high rainfall depth [6], or to high wind velocity values. Figure 4.13 shows that, on a given site, with an invariable land-use and crop management, the long-term average soil loss is dominated by a few and relatively rare events.

Soil erosion is generally a normal aspect of landscape development in which soil particles are removed by wind or water. In some parts of the world, other processes of denudation such as soil mass movement can dominate.

Wind erosion is the process of detachment and transport caused by fluid (air) action on the soil surface [7]. The process removes the finer particles and the organic matter from the top soil. Redeposition of the soil particles can bury soil and vegetation. The process

14 8 12 16 20 24 YEAR

Figure 4.13. Ordered annual soil erosion amounts measured, for a 24-year period, at Kingdom City, MO.

14 8 12 16 20 24 YEAR

Figure 4.13. Ordered annual soil erosion amounts measured, for a 24-year period, at Kingdom City, MO.

operates in a variety of natural environments that lack a protective cover of vegetation. Such areas include the Great Plains of North America, the fringes of arid Africa, India, Australia, and the steppes of Russia, Mongolia, and China.

Soil particles are carried by wind into suspension, by saltation, and by surface creep, depending on their size [1]. Soil particles and small aggregates (<0.05 mm in diameter, 0) are kept in suspension by air turbulence unless the wind velocity is drastically reduced. Intermediate-size grains (0.05 < 0 < 0.5 mm) move in a series of short leaps, jumping into the air and bouncing back on the soil surface. Soil particles larger than 0.5 mm are not lifted. However, grains that are 0.5 < 0 < 1 mm are bumped along the surface by jumping particles.

The wind erosion phenomenon is controlled by soil susceptibility to particle detachment and by the detachment and transport capacity of wind. Factors affecting soil susceptibility to wind erosion are dry aggregate size distribution, mechanical stability of soil structure, surface ridges, rainfall, length of the exposed area, and vegetative cover.

Grains larger than 1 mm in diameter are non-erodible whereas particles that are 0.5 <0 < 1 mm are only eroded by high wind velocities. Soil particles able to move into suspension are highly erodible. Obviously, soil properties such as texture, organic matter, and exchangeable cations, which promote aggregate stability, reduce wind erosion susceptibility. Surface ridges, which increase soil surface roughness, reduce wind velocity near the ground and promote trapping of the eroded particles. Rainfall moistens the soil surface, which transitorily reduces wind erosion. However, rainfall also can promote wind erosion by breaking soil aggregates and smoothing soil surface. Because wind transport capacity at a specific shear velocity u* [m/s] can be considered constant,

Figure 4.14. Displacement of the zero velocity plane due to the vegetation cover. Source: [1].

the distance the wind must travel to reach its load capacity depends on soil erodibility. Vegetative cover is the most effective way to reduce wind erosion because plant cover determines a displacement D + Z0 (m) of the zero plane, in which Z0 is the effective roughness height, that is, the plane at which wind velocity is zero (Fig. 4.14). Plant protection is affected by the amount of cover and time of year in which it is provided, plant geometry, and row orientation. Crop residues left on the soil surface act usefully to reduce wind erosion.

The erosive power of wind is controlled by shear velocity, u* = (r/e)1/2, in which t is the surface shear stress (kg/m2) and e is air density (kg s2/m4); u* is related to the velocity profile and to the drag exerted by wind on the soil surface. For highly turbulent air flow (for shear Reynolds number Re* = u* Z0/v > 90, v being the kinematic viscosity of the fluid), shear velocity is related to the local mean wind velocity uz (m/s) at height z (m) by the logarithmic velocity profile:

in which k = von Karman's constant approximately equal to 0.4. Both the detachment (Dc) and the transport (Tc) capacities of wind depend on u*. In particular, Dc (kg • s-1 m-2) depends on the square of the shear velocity and the size of the erodible particles; Tc (kg • s-1m-2) is essentially proportional to the third power of the shear velocity [1].

For each soil and surface condition, a threshold shear velocity u*c, that is, a minimum wind velocity starting soil particle movement, can be defined. Bagnold [8] and Chepil [9] showed that the critical shear velocity varies with the particle size (Fig. 4.15). In particular, finer particles are characterized by u*c values decreasing for increasing grain size. In fact, the cohesiveness forces are most effective for small soil particles, which also are protected by the surrounding coarser particles. For grain sizes greater than 0.1 mm, u*c increases with the particle diameter because of the increase of the grain weight.

The soil instability process, called "mass movement," usually is neglected in soil erosion studies because this process generally involves high volumes and deep layers of soil.

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Figure 4.15. Relationship between critical shear velocity of wind and particle size. Source: [29].

The instability mechanism depends on the breaking and mass transport processes (breakdown, sliding, rolling, or mixed mechanism). Mass movement occurs as creep, slides, rock falls, debris flow, and mudflow, depending on the ratio between the solid and liquid components of the moving mass. In other words, the different types of mass movement can be considered as part of a continuum of solid transport phenomena ranging from slides, in which the solid/liquid ratio is high, to mudflow having a low solid/liquid ratio.

In 1947, Ellison [10, 11] defined soil erosion as "a process of detachment and transportation of soil materials by erosive agents." For water soil erosion, these agents are rainfall and runoff. Ellison's definition can be extended to take into account deposition processes occurring when the energy of the transporting agent is no longer available to transport soil particles. The intensity of the erosion process depends on the quantity of soil supplied by detachment and the capacity of the erosive agents to transport it. When the agent has the capacity to transport more soil than is supplied by detachment, the erosion is detachment limited. When more soil is supplied than can be transported, the process is transport limited.

According to a classic scheme of the erosive process, the following four phases are distinguished: rainsplash, sheet, rill, and gully erosion.

The impelling force, caused by the raindrops hitting the soil surface, determines soil particle detachment and transport (splash erosion). Waterdrop impact forces depend on the number of hitting drops per unit area and time, drop size distribution, and drop fall velocity. Both rainfall kinetic energy and momentum, which are the most used erosivity parameters, can be calculated using this basic information. The drop size distribution of rainfall is represented briefly by the median (cfeo) drop diameter. Figure 4.16 shows that median drop size increases with rainfall intensity I up to 100 mm/h; at higher intensities,

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Figure 4.16. Relationship between median drop diameter and rainfall intensity. Source: [29].
Allowable Axial Load Coeffient Chart
Figure 4.17. Relationship between fall velocity and drop diameter for a given fall height. Source: [14].

djQ remains essentially constant [12] or decreases [13] because of drop size instability, typical of tropical rainfall. In fact, at higher intensities, drop diameter d is unstable and drop breaks due to both turbulence phenomena and the weight action not counteracted by surface tension. For these reasons, natural drops have a maximum size equal to 67 mm. Drop fall velocity is also strongly dependent on drop size (Fig. 4.17) because the drag force of the waterdrops is contrasted by the gravity force. Figure 4.17 also demonstrates that fall velocity is a function of drop size and fall height; for fall heights greater than 10 m, the fall velocity, called "terminal," depends only on drop diameter

Sharma and Gupta [15] demonstrated that a threshold kinetic energy or momentum exists before the detachment process can be initiated by raindrop impact. The threshold erosivity concept assumes that a minimum energy is needed to overcome the inherent soil strength.

The largest portion of raindrop energy is expended to form an impact crater and to move soil particles. The mechanical breakdown of soil aggregates due to drop impacting can induce a surface seal formation. The most important consequence of seal formation is a reduction of infiltration capacity, which, by increasing surface runoff, can cause an increase in soil erosion. Splash detachment is higher in soils that are not highly susceptible to surface sealing.

Drop impact is more effective if a thin water layer covers the soil surface. This is believed to be due to the turbulence that impacting raindrops induce in the water layer. However, if water depth is higher than a threshold value, ranging from 0.2d to d [16-18], the rainfall energy is dissipated in the water and does not have erosive effects.

Soil detachment by rainfall impact is the main process controlling interrill soil erosion because the detachment capability of sheet flow is negligible compared with that of rainfall because of the low shear stresses of the thin sheet flow [19, 20]. In fact, for soil surfaces, the shear stresses of overland flow are on the order of pascals whereas the soil shear strength is on the order of kilopascals [21].

Soil particles detached by raindrops are encapsulated into the droplets generated after impact and, for sloping surfaces, carried downslope. Transport by rain is generally low and is caused by the component of the raindrop velocity parallel to the surface of the slope.

Rainfall excess occurs on hill slopes when rainfall intensity exceeds soil infiltration capacity. According to the classic Horton scheme [22], at the top of the hill slope, a flowless zone occurs. Flow begins at a critical distance downslope from the divide. Farther downslope, water depth increases and flow becomes channeled and breaks up into rills. The field-scale runoff process is characterized by rainfall excess dominated runoff occurring as shallow sheet flow or flow in small concentrated channels. The runoff response to rainfall is basically controlled by rainfall intensity and soil properties. For modeling purposes, the rainfall excess approach uses a time-intensity rainfall distribution and an infiltration equation, such as the Chu's [23, 24], to compute a rainfall excess distribution over the field. For field-scale applications, some form of the De Saint Venant shallow water equations has been used recently to route the rainfall excess over the flow surface [25].

Moss [26] showed that overland-flow sediment transport is a combined action of raindrops and flow: the raindrop impact induces the flow to transport particles locally increasing its turbulence. In other words, without rainfall, the flow would be incapable of transportation. Flow transport processes associated with raindrop action are called rain-induced flow transport (RIFT) [27, 28]. RIFT acts for shallow flow depth (less than 3d) impacted by medium to large-size raindrops [21].

The hill-slope flow rarely is distributed evenly on the soil surface. More commonly, it appears as a mass of anastomosing water courses with no pronounced channels [29]. Rills are ephemeral features, that is, small and intermittent water channels that do not interfere with conventional tillage operations. Once obliterated, rills will not reform at the same site [30]. Merritt [31] identified four subsequent stages of rill development: sheet flow, flow lines (starting of flow concentration), microchannels without headcuts, and microchannels with headcuts (channeled flow).

Compared with rill erosion, interrill erosion contributes a very small proportion to the sediment transported downward [20].

Rill initiation usually is described by the threshold value of a variable crucial for this erosion mechanism. Generally, a hydraulic variable, such as discharge [32], Froude number [33], shear velocity [34], or unit stream power [35], is used to describe the ability of the erosive agent to start rilling. Other authors suggest the consideration of the soil susceptibility to rilling. For example, Savat [36] explained rilling by defining a critical Froude number Fc depending on the median soil grain size. A more complex approach, proposed by Boon and Savat [37], includes both Fc and the sediment concentration in rill flow.

For a recently tilled field, rill initiation also can be induced by piping. When the topsoil is saturated, in some isolated location an unequal settlement of the surface layer can take place. This phenomenon may be due to the large pores among clods. Runoff from the upper area flows into the crevices, resulting from the unequal settlement, and creates pipes just above the undisturbed, more compact subsoil [35]. The main factors controlling piping are soil properties, such as porosity and soil erodibility.

Since rill discharge significantly affects the ability of the rill to detach and transport sediment particles, the knowledge of the number of rills that may form per unit of cross-sectional area and the variation in flow rate between individual rills is necessary. Gilley et al. [38] studied partitioning flow between rills and determined the relative frequency of flow rates among rills on a given plot. Figure 4.18 shows the distribution of the relative discharge, equal to the discharge for each rill divided by the maximum discharge on the plot among rills. The figure shows that differences in discharge existed between individual rills and that 30% of the rills had discharge equal to the maximum value. In

Figure 4.18. Relative frequency of measured relative discharge. Source: [38].

addition to discharge, identification of other rill hydraulic variables, such as rill width, hydraulic roughness coefficient, and flow velocity, also may be important [39, 40].

Rill erosion is the detachment and transport of soil particles by concentrated flow. Soil particle detachment by flow depends on the rill erodibility, the hydraulic characteristics of rill flow, and the actual flow sediment load. The simplest approach is considering rill detachment as due only to the scouring processes of the wetted perimeter. The maximum rill detachment, called detachment capacity, Dc (kg ■ s-1m-2), occurs when a clear flow moves on an erodible soil. The soil is characterized by a rill erodibility parameter, Kr, lumping the effects of different factors such as grain size distribution, rock fragment cover [41], and soil structure and its stability. For a given soil, the detachment capacity Dc depends on the excess of flow energy content as related to a threshold value. The most widely applied equation to estimate Dc is the modified Du Boys sediment transport equation [42], in which t is the bed shear stress (Pa), tc is its threshold value, and a is a constant quasi-equal to 1 [43, 44]. Other approaches assume as flow variable the discharge or the stream power [45]. A more detailed approach for estimating Dc needs to take into account scouring, headcutting [46], side-wall sloughing, and slaking [47]. Kohl [48] found that head-cutting accounted for up to 60% of total rill erosion on some of the soils considered during Water Erosion Prediction Project development. Flow stream power is used as an indicator of detachment due to headcutting [47]. Side-wall sloughing could be a major erosion component in freshly tilled soils with low cohesion and high capillary pressures that have a rill caused by scour or headcut erosion [49]. Slaking affects Dc only for soils with high clay content, low organic matter, and low antecedent water content [50] or soils with a weak structural stability [51]. Establishing the influence of the above-mentioned factors on Dc needs experimental evaluation of soil erodibility parameters corresponding to each process. Soil structure mostly affects the values of these erodibility parameters [47].

Since sediment generally is carried by runoff water, the actual detachment rate Dr (kg ■ s-1m-2) is less than Dc. According to Foster and Meyer [52], the detachment capacity has to be reduced by a feedback factor fc that depends on the ratio between the actual sediment transport G (kg ■ s-1m-1) and the transport capacity Tc (kg ■ s-1m-1):

Tc expresses the maximum sediment discharge that can be transported by a rill flow with given hydraulic conditions. Tc generally is assumed to be proportional to the 1.5 power of the bed shear stress [53]. The feedback factor expresses the physical circumstance that the rill flow has to detach the sediment amount necessary to make the difference between Tc and the actual sediment load negligible. From an energy point of view, the flow energy available for rill detachment is less than the total flow energy because a quota is expended to carry the actual suspended sediment load G. When the sediment transport capacity is exceeded by the sediment load, deposition occurs. For small channels, such as rills, the deposition rate is assumed to be proportional to the difference between the actual

sediment load and the transport capacity. Foster [20] assumed that the proportionality constant was directly proportional to the settling velocity and inversely proportional to the discharge.

Gullies are relatively permanent, steep-sided channels in which ephemeral flows occur during rainstorms [29] and cannot be eliminated by usual tillage operations. Gullies are usually deep channels with a narrow cross section. In the first stage, the gully cross section is V-shaped. As the gully develops, its cross section can be modeled by scouring and side-sliding phenomena for assuming a triangular, trapezoidal, and U-shape. The gully channel is characterized by an overfall at the gully head, advancing upstream [54].

The initiation and growth of gullies are dependent on a flow concentration sufficient to form a definable channel. Schumm [55] suggested that the channel length is dependent on the contributing drainage area.

According to Mitchell and Bubenzer [54], gullies are formed when rills combine and develop to the extent that they cannot be eliminated by tillage operations. Morgan [29] established that gully initiation is a more complex process. Small depressions on the hill slope, for example due to a break in vegetation cover (Fig. 4.19), determine flow concentration inducing localized erosion processes. In particular, the erosion is concentrated at the head of the depression where a near-vertical scarp develops. Water falling from the upstream hill slope into the depression determines scouring at the base of the headcut leading to collapse and retreat of the scarp upslope. Flow concentration induces gully floor incision and the development of a stable channel by the scouring action of a running channel flow.

Gully development is not always due to surface erosion processes. In fact, concentrated runoff occurring as subsurface pipe flow can determine erosion processes giving rise to the development of a subsurface tunnel network. Heavy rain can induce subsidence of the ground surface, so exposing the pipe network as gullies.

Haigh [56] described for a desert gully system a complex mechanism of gully enlargement (Fig. 4.20a) caused by both scouring surface processes and tunnel erosion. According to this scheme, an increase of gully cross-sectional area is due to a parallel retreat of gully walls, an aggradation of gully bed, and an enlargement of soil pipe by collapse (Fig. 4.20b). Pipe breaks the gully bed, creating a narrow, vertical sided slot (Fig. 4.20c) inducing a parallel retreat of the former soil pipe and aggradation of the channel bed (Fig. 4.20d).

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