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a Dep. of Soil Science, Egerton University, P.O. Box 536, Njoro, Kenya
b Institute of Soil, Water and Environmental Sciences, The Volcani Center, Agricultural Research Organization, P.O. Box 6, Bet Dagan, 50250, Israel
* Corresponding author (meni{at}agri.gov.il)
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ABSTRACT |
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 TOPNOTES ABSTRACT INTRODUCTIONMATERIALS AND METHODSRESULTS AND DISCUSSIONSUMMARY AND CONCLUSIONSREFERENCES |
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9.3 mm h-1 in the remaining soils. Scanning electron microscope (SEM) observations indicated that the kaolinitic soil had a thin crust (
0.1 mm) containing large particles (0.1 mm), whereas the montmorillonitic soils had thicker crusts (>0.2 mm) comprising either small (0.02 mm) particles with a very developed washed-in zone underneath or large (0.2 mm) ones with fine material between them. The crust layer in the nonphyllosilicate soils was 0.2 mm thick and composed of fine particles 0.01 mm. The high aggregate stability and the low dispersivity of the kaolinitic soil, which minimized soil detachment, and its low runoff, which minimized its transport capacity, limited the soil loss to 0.33 kg m-2, whereas the low aggregate stability and high runoff of the montmorillonitic soils contributed to their soil losses of 1.24 and 1.14 kg m-2. The intermediate aggregate stability and the high runoff of the nonphyllosilicate soils accounted for their intermediate soil losses of 0.75 and 0.8 kg m-2.
Abbreviations: CEC, cation-exchange capacity • EC, electrolyte conductivity • ESP, exchangeable Na percentage • IR, infiltration rate • MWD, mean weight diameter • SEM, scanning electron microscope
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INTRODUCTION |
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TOPNOTESABSTRACTINTRODUCTIONMATERIALS AND METHODSRESULTS AND DISCUSSIONSUMMARY AND CONCLUSIONSREFERENCES |
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McIntyre (1958) described the general sequence of events that leads to crust formation under rainfall conditions as follows: (i) breakdown of soil aggregates caused by raindrop impact or slaking; (ii) movement of fine particles into the upper few millimeters of the soil, and deposition in the voids; and (iii) compaction of the soil surface to form a thin film, which restricts further entry of water and movement of soil particles.
Two main types of soil crust, namely structural and depositional crusts, are generally recognized, according to their mechanisms of formation (Chen et al., 1980). Structural crusts are due mainly to water-drop impact, whereas depositional crusts are formed by translocation of fine particles and their deposition at some distance from their original location. In contrast, Valentin and Bresson (1992) classified crusts into three main groups: structural, depositional, and erosional crusts, of which the last are formed through erosion of the sand layer at the top of the structural crust, following runoff initiation.
Luk et al. (1990) subjected some loess soils, dominated by illite and kaolinite minerals, to rainstorms of a range of durations. Scanning electron microscope observations of the crusts in these soils revealed a rearrangement of the microfabric at the soil surface. Gal et al. (1984) and McIntyre (1958) reported that in the final stage of crust formation, the crust comprised a 0.1-mm-thick skin composed almost entirely of fine particles, and a 2-mm-thick washed-in layer. Moreover, Chen et al. (1980) observed a crust on a sandy loam soil, that comprised coarse particles stripped of the fine ones.
The tendency of a soil to form a crust depends on aggregate stability (Le Bissonnais, 1996). Goldenberg et al. (1988) classified the factors that affect aggregate stability as solution factors and soil factors. Many studies have addressed only solution factors, including: electrolyte conductivity (EC) (Shainberg and Letey, 1984), pH (Suarez et al., 1984), and soluble silica (Shanmuganathan and Oades, 1983). Soil factors that have received attention include: exchangeable Na percentage (ESP) (Levy and van der Watt, 1988; Ben-Hur et al., 1998), organic matter content (Le Bissonnais and Arrouays, 1997), and texture and carbonate content (Ben-Hur et al., 1985).
Aggregate stability has been found to increase with increasing clay content (Kemper and Koch, 1966) but, conversely, Moldenhauer and Kemper (1969) found that increasing clay content promoted crust formation. Ben-Hur et al. (1985) explained this paradox by suggesting that in soils containing >20% clay, the clay fraction acts as a cementing material, stabilizing soil aggregates against the beating action of raindrops, and so preventing crust formation. On the other hand, in soils containing 20% clay, the clay acts as a substrate for crust formation, decreasing the steady-state hydraulic conductivity of the crust.
Soil mineralogy has a substantial effect on aggregate stability and dispersion. Singer (1994) reviewed the effects of clay mineralogy on soil dispersivity and concluded that smectitic soils are the most dispersive and kaolinitic soils the least dispersive. The dispersivity of illitic soils is intermediate, but may sometimes exceed that of smectitic soils. In soils dominated by 2:1 clays, the aggregate stability is affected mainly by polyvalent metal-organic matter complexes that form bridges between the negatively charged clay platelets (Six et al., 2000). However, the stability in 1:1 clay-dominated soils is attributed mainly to the binding capacity of the minerals themselves (Oades and Waters, 1991). Soils that are dominated by clay-sized minerals, such as quartz and feldspar, disaggregate preferentially over soils containing kaolinite and smectite when these soils are shaken in distilled water (Neaman et al., 1999; Bühmann et al., 1996).
However, in spite of the dominant effect of soil mineralogy on soil dispersivity and aggregate stability, little is known about its effects on seal formation and micromorphology, runoff and soil loss. The effects of soil mineralogy on crust formation IR, and erosion were studied for some soils from South Africa and Israel by Stern et al. (1991), who used a laboratory rainfall simulator; they suggested that soils in which either kaolinite or illite predominated, but which contained small amounts of smectite, were as susceptible to crust formation as smectitic soils. However, smectitic soils were much more erodible than soils that contained only small amounts of smectite, and the soils that did not contain smectite were the least erodible. Other investigators have reported similar findings on erosion (Ben-Hur et al., 1992) and final IR (Ben-Hur et al., 1992; Levy and van der Watt, 1988), for a range of soils with various textures and organic matter contents.
In a laboratory rainfall simulator study of selected loess soils with various properties, Romkens et al. (1995) found that the presence of highly expansive smectite clay in the soil caused a rapid reduction in IR despite the high organic C content and the coarse texture of the soil; this indicated the importance of soil mineralogical constituents for crust development. A similar conclusion was reached by Mermut et al. (1997) who found that the amount of splash was about four times greater with a loamy soil, in which smectite, mica, and vermiculite were the dominant clays, than with a silt loam soil, in which vermiculite, mica, and kaolinite dominated. The objective of the present study was to examine the effects of soil mineralogy and texture on crust micromorphology, infiltration, runoff, and soil loss. Specifically, aggregate stability, crust thickness and structure, IR, runoff, and erosion in three groups of soils with differing mineralogies were determined.
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MATERIALS AND METHODS |
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TOPNOTESABSTRACTINTRODUCTIONMATERIALS AND METHODSRESULTS AND DISCUSSIONSUMMARY AND CONCLUSIONSREFERENCES |
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x-ray source (Whittig, 1965). Estimates of the amounts of the various clay minerals were derived from the relative peak areas on the diffractograms.
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Fast-wetting Aggregate Stability Test.
The soil aggregate stability was determined by the fast-wetting method proposed by Le Bissonnais (1996). This method was chosen because in the rainfall simulator study, which is described below, the soils were under fast-wetting conditions (Levy et al., 1997). Soil samples with aggregates of 3 to 4 mm were put in the oven at 40°C for 24 h, and 5 g of the oven-dry aggregates were gently immersed in a beaker containing 50 cm3 of deionized water for 10 min. The water was then sucked off with a pipette. The soil material was transferred to a 50-µm sieve that had previously been immersed in ethanol and gently moved up and down in ethanol five times to separate the <50-µm fragments from the >50-µm ones. The >50-µm fraction was oven-dried and then gently dry sieved by hand on a column of sieves of mesh sizes 2, 1, 0.5, 0.25, 0.1, and 0.05 mm. The weight of each fraction was then calculated; that of the <50-µm fraction was the difference between the initial weight and the sum of the weights of the other six fractions. The aggregate stability for each soil sample was expressed by calculating the mean weight diameter (MWD) of the seven classes:
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i. Rainfall Simulator Study.
Disturbed soil samples were air dried, crushed to pass through a 4-mm sieve and mixed thoroughly. To determine the aggregate-size distribution in the soil samples, a separate 40-g sample of each soil was subjected to dry sieving through a set of sieves with 2-, 1-, 0.5-, 0.25-, 0.1-, and 0.05-mm openings. The soil samples, in triplicate, were then packed in 0.5 by 0.3 m perforated trays in an 0.02-m-thick layer, to a mean bulk density of 1.29 (±0.03) Mg m-3. The trays were then placed over a 0.08-m-thick layer of coarse sand on a 9% slope in a rainfall simulator of the rotating disk type (Morin et al., 1967). The samples were first prewetted from below with tap water (EC = 0.7 dS m-1) and then subjected to an 80-mm rainstorm of deionized water (simulated rainwater quality). The mean diameter of the raindrops was 1.9 mm, median drop velocity 6.02 m s-1, kinetic energy 18.1 J mm-1 m-2, and rain intensity 40 mm h-1.
During the rainstorm, the volume of rainfall percolating through the soil was measured and the IR was calculated. From the moment that runoff was observed to commence, it was collected until the end of the rainstorm. The runoff samples from each replicate were analyzed by a wet-sieving procedure similar to that described by Gabriels and Moldenhauer (1978), to determine the particle-size distribution of the eroded sediment. Each sample was gently poured in bulk through a nested set of 2-, 1-, 0.5-, 0.25-, 0.1-, and 0.05-mm sieves. Extreme care was taken to ensure a high degree of uniformity in sieving operations and to avoid creating artifacts. The runoff that contained sediment of particle size <0.05 mm was gently put into 1-L cylinders and the clay concentration was determined with a hydrometer.
Scanning Electron Microscope Observations.
The trays with the soils that were crusted after the rainstorm were left to dry for a week. A piece of crust was carefully broken along its natural planes of weakness into 15-mm fragments. The top of the crust was mechanically stabilized by coating it with a gold layer. The sample was glued onto the SEM stub with conductive carbon glue and the fracture was covered by sputtering with a thin layer of gold, 40 nm in thickness. The samples were then placed in the SEM and micrographs were taken.
Statistical Analysis.
All the studies were conducted in three replicates, and the data were subjected to analysis of variance as a complete randomized design (Steel and Torrie, 1981). Separation of means was tested according to Tukey's honestly significant difference, at the 0.05 significance level.
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RESULTS AND DISCUSSION |
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TOPNOTESABSTRACTINTRODUCTIONMATERIALS AND METHODSRESULTS AND DISCUSSIONSUMMARY AND CONCLUSIONSREFERENCES |
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9.3 mm h-1) and significantly different from that of the Tunyai soil. The final IR and the total runoff values shown in Fig. 3 are not consistent with the MWD values (Table 2), that divided the soils into three groups, according to their aggregate stability. This indicates that the aggregate stability of the soil is not the only factor that affects crust formation and IR. The rearrangement of particles in the crust, and its thickness should affect the IR of the soil under sealing conditions.
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0.1 mm) of slightly compacted aggregates 0.1 mm in size. In contrast, beneath this 0.1-mm uppermost layer, large aggregates, >0.3 mm in size, were observed. The structure of the layer immediately under the uppermost crust layer (Fig. 4A) was similar to that of the bulk soil at 5-mm depth, which was not affected by the rainfall (Fig. 4B). This suggests that the soil layer immediately below the crust layer (Fig. 4A) was not disturbed by raindrop impact, and no washed-in layer is apparent on the SEM micrograph of this layer. This crust characteristic limited the sharp reduction of the hydraulic conductivity of the seal and the soil IR (Fig. 3).
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0.2 mm in thickness (skin layer) was observed (Fig. 4A), which was packed with aggregates 0.02 mm in size. Under this layer a very dense and compacted layer 0.5 mm thick was formed; this was a washed-in zone, in which dispersed clay particles, transferred from the upper layer with the infiltrated water, accumulated (Fig. 4B). The high clay content of the Neve Ya'ar soil, and the high dispersivity of montmorillonite, which is the dominant clay in this soil, allowed the high accumulation of the clay particles in this washed-in zone.
In the Molo soil an upper layer, 0.2 mm in thickness, comprising well-compacted particles 0.01 mm in size, was observed (Fig. 4A). This upper layer was the crust, and its structure was significantly different from that of the bulk soil at the 5-mm depth (Fig. 4B). In the lower soil layer (Fig. 4B) relatively large aggregates, measuring 0.3 mm with small microaggregates between them were observed.
In the Netanya soil, the structure of the upper 3-mm layer—the crust layer—was different from that of the bulk soil (Fig. 5). In the crust layer, partly naked sand particles 0.2 mm in size, with fine material between them, were observed. In contrast, in the bulk soil, the sand particles were coated with fine materials, and there was almost no fine material in the pores between the sand grains. The crust layer showed two zones: an uppermost 0.5-mm layer with higher density, which was underlain by a 2.5-mm layer with lower density. As discussed by Bielders et al. (1996) for coarse-textured soil, the high density of the 0.5-mm uppermost layer of the crust could result from infilling of pores by fine material and compaction by the raindrop impacts. It could be that this topmost layer is the clay skin of an erosion crust and that the relative accumulated fine material between 0.5- and 3-mm depth was the washed-in layer. However, because of the low clay content in Netanya soil, the amount of dispersed material was small, therefore, there was only limited accumulation of fine materials in the washed-in layer.
The low final IR values (9.3 mm h-1) of the Neve Ya'ar, Netanya, Molo, and Njoro soils (Fig. 3) can be explained in terms of the dense structures and relatively large thicknesses (>0.2 mm) of the crusts that developed in these soils (Fig. 4 and 5). These crusts comprised either only small particles (0.01 mm), or an uppermost layer of particles measuring 0.02 mm with a highly developed washed-in zone beneath it, or relatively large particles (0.2 mm) with fine material between them. In contrast, in the Tunyai soil, characterized by a high final IR (20.5 mm h-1) (Fig. 3), the crust was thin (0.1 mm) and it contained relatively large particles (0.1 mm in size) (Fig. 4A).
Total soil losses for the various soils during the entire rainstorm are presented in Table 2. The soil loss was significantly lowest in the kaolinitic soil (Tunyai), highest in the montmorillonitic soils (Neve Ya'ar and Netanya) and intermediate in the nonphyllosilicate soils (Molo and Njoro). These soil losses were not consistent with the IR values of the soils (Fig. 3), since the soils were divided into only two groups according to their IRs.
Interrill erosion involves two major processes: (i) detachment of soil material from the soil surface by raindrop impact; and (ii) transport of the resulting sediment by runoff flow. Soil detachment depends mainly on rainfall characteristics and the aggregate stability of the soil. In the present study, all the soils were subjected to rainfall with the same characteristics, therefore, the soil detachment depended mainly on the aggregate stability of the soil. The transport of the sediments depends on the runoff characteristics and the particle-size distribution of the impacted soil surface. Because in the present study all the soils were subjected to rainfall onto a 9% slope, the main runoff characteristic which determined the runoff transport capacity was the runoff rate. Under these conditions, the soil loss values of the various soils can be explained as follows. The soil loss in the kaolinitic soil was the lowest (Table 2) because the aggregate stability of this soil was the highest (high MWD value in the fast-wetting test, Table 2), which decreased the soil detachment, and because the runoff rate obtained during the rainstorm in this soil was the lowest (Fig. 3), which decreased the runoff transport capacity. The montmorillonitic soils, with the lowest aggregate stability (Table 2) and the highest runoff rate (Fig. 3) showed the highest soil loss (Table 2). In the case of the nonphyllosilicate soils, their intermediate aggregate stability (Table 2) and their high runoff rates (Fig. 3) contributed to the intermediate soil losses found in these soils (Table 2).
The particle-size distributions of the sediments in the runoff and their MWD values for each soil are presented in Fig. 6
and Table 2, respectively. The particle-size distribution in the runoff was controlled mainly by that at the impacted soil surface and by the runoff transport capacity during the rainstorm. The high clay content in the montmorillonitic Neve Ya'ar soil and the high dispersivity of montmorillonite probably led to the high frequency of very small particles at the impacted soil surface. This led to the high frequency (80%) of clay-size particles (Fig. 6) and the very small (0.03 mm) MWD in the runoff from this soil (Table 2). In contrast, in the montmorillonitic Netanya soil, which contained 90% sand and 10% clay (Table 1), the sand particles apparently had the highest frequency at the impacted soil surface. In this case, the sand particles measuring 0.1 to 0.25 mm in size had the highest frequency (50%) in the sediments (Fig. 6) and the MWD in the runoff was large (0.2 mm) (Table 2). The clay contents in the kaolinitic Tunyai soil and the montmorillonitic Neve Ya'ar soil were similar. However, the MWD in the runoff from the former soil was significantly larger than that from the latter (Table 2). The low dispersivity of the kaolinite in Tunyai soil probably led to a high frequency of relatively large particles at the impacted soil surface and in the runoff (Fig. 6), and to the large MWD (0.12 mm) of the sediments (Table 2). The clay contents of the nonphyllosilicate soils were relatively high (Table 1). Likewise, the particles at the crust in these soils were very small (0.01 mm) (Fig. 4A), indicating that the aggregates at the surfaces of these soils were broken down into small particles. However, the MWDs in the runoff from these soils were relatively large, similar to that from the Netanya soil (Table 2). This suggests that in these soils, some of the aggregates at the soil surface were not broken down or were broken down to relatively large particles that were eroded with the runoff during the rainstorm.
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SUMMARY AND CONCLUSIONS |
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TOPNOTESABSTRACTINTRODUCTIONMATERIALS AND METHODSRESULTS AND DISCUSSIONSUMMARY AND CONCLUSIONSREFERENCES |
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For all the studied soils, the IR decreased with increasing cumulative rainfall until a final IR was reached. This decrease of the IR was a result of the crust formation at the soil surface. The soils were separated into two groups on the basis of their IR values: in the first, comprising the kaolinitic soil, the IR decreased gradually and the final IR was high (20.5 mm h-1); in the second group, comprising the rest of the soils, the IR decreased sharply and the final IR was low (9.3 mm h-1). The aggregate stability of the soils was not the only factor that affected crust formation and IR; the rearrangement of particles in the crust, and its thickness also affected the IR of the soils under sealing conditions.
The low final IR values of the montmorillonitic and the nonphyllosilicate soils can be attributed to the dense structures and relatively large thicknesses (>0.2 mm) of the crusts that developed on these soils. These crusts comprised either only small particles (0.01 mm), or an uppermost layer of particles measuring 0.02 mm with a highly developed washed-in zone beneath it, or relatively large particles (0.2 mm) with fine material between them. In contrast, in the kaolinitic soil characterized by a high final IR, the crust was thin (0.1 mm) and it contained relatively large particles (0.1 mm in size).
The soil loss was significantly lowest in the kaolinitic soil, highest in the montmorillonitic soils, and intermediate in the nonphyllosilicate soils. The kaolinitic soil showed the lowest soil loss because of its high aggregate stability, which decreased the soil detachment, and the low runoff from this soil, which decreased the runoff transport capacity. The montmorillonitic soils, which had the lowest aggregate stability and the highest runoff, showed the greatest soil loss. In the case of the nonphyllosilicate soils, their intermediate aggregate stability and their high runoff rates accounted for the intermediate soil losses from these soils.
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ACKNOWLEDGMENTS |
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NOTES |
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TOPNOTESABSTRACTINTRODUCTIONMATERIALS AND METHODSRESULTS AND DISCUSSIONSUMMARY AND CONCLUSIONSREFERENCES |
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Received for publication March 2, 2001.
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TOPNOTESABSTRACTINTRODUCTIONMATERIALS AND METHODSRESULTS AND DISCUSSIONSUMMARY AND CONCLUSIONSREFERENCES |
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