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DIVISION S-1-SOIL PHYSICS

Salinity and Compaction Effects on Soil Water Evaporation and Water and Solute Distributions

^a Alexandria Univ., Faculty of Agriculture-Damanhoar, Egypt
^b Dep. of Agronomy, Iowa State Univ., Ames, IA 50011 USA

rhorton{at}iastate.edu

	ABSTRACT

TOP NOTES ABSTRACT INTRODUCTION Materials and methods Results and discussion Conclusion REFERENCES

 
Water evaporation and solute transport were studied in open soil columns. The study included two different soil materials — Clarinda clay (fine, montmorillonitic, mesic, sloping Typic Argiaquoll) and Fayette silty clay loam (fine-silty, mixed, mesic Typic Hapludalf) — and three conditions. Two conditions were noncompacted solute-free and salinized noncompacted soil columns of both Clarinda and Fayette soils, and one condition was compacted salinized soil columns of Clarinda soil only. The initial soil water contents were 0.271 and 0.181 m³ m^-3 for noncompacted Clarinda and Fayette soils, respectively. The initial soil water content of compacted Clarinda was 0.393 m³ m^-3. The initial KCl concentrations were 1.11 and 0.92 mol kg^-1 of soil solution for Clarinda and Fayette soils, respectively. Measured ratios of evaporation loss from the noncompacted salinized soil columns to the amount of water evaporated from noncompacted solute-free soil columns increased with time from 0.78 to 0.89 for Clarinda and 0.90 to 0.95 for Fayette soils. Evaporation from noncompacted Clarinda soil increased with time from 0.73 to 0.77 of the evaporation from compacted Clarinda soil. A numerical model of heat, water, and solute transfer was used to predict distributions of temperature, water content, and solute concentration for a given evaporation rate. Efficiency of the model for reproducing the water content and solute concentration ranged from 94.5 to 61.1%. The predicted and observed solute concentrations increased with time in the upper 0.02 m of noncompacted soil. Also, the upper soil portion, the 0.05-m layer, dried drastically. Both observations and predictions indicate complex interactions between heat, water, and chemicals near evaporating surfaces.

	INTRODUCTION

TOP NOTES ABSTRACT INTRODUCTION Materials and methods Results and discussion Conclusion REFERENCES

 
THE EVAPORATION RATE OF WATER from a wet soil is a function of local meteorological (e.g., wind, radiation, air humidity) and soil conditions (e.g., water content, soil texture, salinity, tillage practices). Soil water and salinity redistributions are greatly affected by these environmental conditions. Understanding mechanisms of water and chemical movement in soil under evaporative conditions is helpful in managing soil salinity and water content. Several researchers have focused on soil water evaporation. Gardner (1959) developed a simple theory to describe the evaporation rate from soil columns as a function of isothermal water diffusivity. He stated that the rate of evaporative drying was controlled largely by the surface boundary conditions and soil physical characteristics. Fritton et al. (1967) observed water and Cl^- distributions in laboratory soil columns under evaporation potentials ranging from 0.0051 to 0.0192 m d^-1. They reported that evaporation-zone thickness was a function of the evaporation potential and the time duration of evaporation. The evaporation zones were 0.015, 0.025, 0.037, and 0.06 m deep after 25 d of evaporation potentials of 0.0051, 0.0073, 0.0113, and 0.0192 m d^-1, respectively. The evaporation zones under evaporation potential of 0.01 m d^-1 were 0.01, 0.02, and 0.03 m deep after 7, 13, and 18 d of evaporation, respectively. Hanks et al. (1967) reported that thermal water transfer amounted to 10% of the net upward water flow in soil columns. They reported that for saturated and near-saturated soil conditions, isothermal water flow was sufficient to describe surface evaporation. Similar results considering the thermal effect on water evaporation were reported by Hammel et al. (1981), Scanlon and Milly (1994), and Bendz and Bengtsson (1996).

In addition to the thermal effect on water flow, the solution osmotic potential can be a factor affecting water transfer in soil (Scotter and Raats, 1970; Globus, 1983; Nassar and Horton, 1989; Bear and Gilman, 1995; Salhotra et al., 1985; Shimojima et al., 1996). Evaporation from salinized and fresh water surfaces was studied by Salhotra et al. (1985). They reported that the ratio of evaporation from a salinized water body to evaporation from a fresh water body ranged from 0.94 (Mediterranean Sea water with a density of 1037 kg m^-3) to 0.69 (Dead Sea water with a density of 1233 kg m^-3). This reduction of evaporation from the salinized water depends on the saturation vapor pressure and the ionic composition of the salinized water. Similarly, the presence of a large solute concentration in soil solution can reduce soil water evaporation. Solution concentration can also affect the evaporation rate through its effect on soil hydraulic properties (Lima et al., 1990).

Compaction can be a major factor affecting soil hydraulic properties and soil water flow. Warkentin (1971) reported that compaction increased plant-available retention of water for both clay and sand soils. The magnitude of increased water retention became smaller as soil water matric potential decreased, that is, the effect of bulk density was less at -150 m than at -1 m of soil water matric potential. Affleck (1980) investigated water flow in layered soil columns. The columns had high bulk density in the top layer and low bulk density in the lower layer. He found that upward moisture flows through the columns increased with increasing bulk density of the top layer.

Our study had two objectives. The first objective was to observe the effect of salinity (two levels) and soil compaction (two levels) on water evaporation, water content distributions, and solute concentration distributions in soil. The second objective was to use a theory of energy and mass transfer to predict soil water and solute concentration distribution in response to surface evaporation under salinity and compaction conditions.

	Materials and methods

TOP NOTES ABSTRACT INTRODUCTION Materials and methods Results and discussion Conclusion REFERENCES

 
Experiment
Subsurface soil materials were collected from two Iowa soils, a loess and a paleosol. The loess-derived soil was a Fayette (fine-silty, mixed, mesic Typic Hapludalf). The exhumed paleosol was a Clarinda (fine, montmorillonitic, mesic, sloping Typic Argiaquoll). The textures of Fayette and Clarinda soils are silty clay loam and clay, respectively. Selected soil physical properties of these soil types are reported in Horton et al. (1987). The soil materials were collected by excavation, air dried, and ground to pass a 1-mm sieve. The soil was moistened to a desired initial soil water content by adding and mixing either KCl solutions (salinized batch) or distilled water (solute-free batch). The moistened soil was covered and stored at 20°C for more than 4 d, then packed into polyvinyl chloride columns (0.04 m in diameter and 0.20 m long). The initial gravimetric water contents were 0.181 and 0.271 kg kg^-1 for Fayette and Clarinda soils, respectively, for both batches. The initial solute concentrations for the salinized batch were 0.92 and 1.11 mol kg^-1 of soil solution for the Fayette and Clarinda soils, respectively. For the compacted (salinized) treatment, Clarinda soil was packed to a bulk density of 1.45 Mg m^-3. For all other treatments, Clarinda or Fayette soil was packed to a bulk density of 1.0 Mg m^-3. The soil columns packed at low density were called noncompacted soil columns. Soil water retention curves for noncompacted and compacted soils were determined by pressure plate extraction and thermocouple psychrometer chamber equilibrium. The initial soil water matric potentials for the noncompacted soil columns were estimated by interpolation from the retention curve data (Campbell, 1974) to be -24 and -25 m for the Fayette and Clarinda soils, respectively. Likewise, the initial soil water matric potential of the compacted Clarinda soil was approximately -0.9 m. A mechanical device acting as a controlled hydraulic press (Instron Model 1125, Instron Corp., Canton, MA) was used to obtain the desired density within the compacted soil columns. Columns were packed in 0.05-m increments in an attempt to provide uniform compaction along the whole column length. Each soil column was closed at the bottom end and left open at the top end. The soil columns were buried vertically in a greenhouse soil pit. All column tops were level with the bed soil surface. The columns were exposed to natural radiation for 28 d. One thermocouple (copper-constantan) was placed at each end of the soil columns (0.0 and 0.20 m) to measure boundary temperatures. Temperature was measured hourly and recorded with a CR5 digital recorder (Campbell Scientific, Inc., Logan, UT). Fourier series were fitted to the measured diurnal boundary temperatures. An example of the fitted boundary temperature is presented in Fig. 1 . Surface diurnal temperature amplitudes ranged from 5.0 to 9.0°C during the course of the experiment. Soil columns were weighed daily for the first 8 d and every 2 to 3 d thereafter to determine cumulative evaporation. Soil columns were weighed between 0700 and 0800 h when the soil temperature gradients were relatively small. At two different times, 14 and 28 d, three soil columns of each treatment were removed and sectioned. Soil water content was determined gravimetrically in each soil section. From measures of electrical conductivity, solute concentrations were estimated in 1:5 soil/water extracts. In summary, 24 soil columns were used, including two soil types, two salinity levels, two sampling dates, and three replicates. In addition, three columns were used for salinized compacted Clarinda soil at sampling date of 14 d.

View larger version (22K): [in this window] [in a new window]   Fig. 1 Fitted soil temperatures (solid and dashed lines) at 0.00- and 0.20-m depths using Eq. [5] and [6], respectively and measured temperature (symbols)

(1)

where q_h is the net heat flux (W m^-2); the coefficients K₁, K₂, and K₃ are storage terms for heat; t is time (s); z is depth (m); and (m³ m^-3), T (°C), and C (mol kg^-1 of soil solution) are water content, temperature, and solute concentration, respectively.

The nonsteady-state equation for soil water flow may be written as

(2)

where q_v and q_L are the mass fluxes of vapor and soil water, respectively (kg m^-2 s^-1); _w and _L are the density of pure water and soil water, respectively (kg m^-3); and C₁, C₂, and C₃ are storage terms for water.

The nonsteady-state equation for solute transport may be written as

(3)

where n_c is the net flux of solute (mol m^-2 s^-1). Nassar and Horton (1997) stated that solute moves by dispersion–diffusion, salt sieving, and salt thermal diffusion. Nassar et al. (1997a) found that thermal diffusion was not significant under cyclical surface temperature conditions. Accordingly, the thermal diffusion of solute was neglected in the present study. During evaporation of water and redistribution of the solute, precipitation of the solute can occur if the solute concentration reaches the solubility limit. So, the precipitation was taken into account to predict solution osmotic potential and solute transport in soil (Nassar et al., 1997b).

The initial conditions associated with the heat and mass transfer equations, Eq. [1], [2], and [3], are

(4)

where L is soil column length (m).

The diurnal temperature boundary conditions at both ends, T(0,t) and T(L,t) can be described, respectively, by

(5)

(6)

where T₀, T_L, A, B, and are Fourier constants, and is diurnal angular frequency [rad s^-1]. Equations [4] and [5] were fitted to the measured temperatures for each day at the upper and lower boundaries, respectively, so the Fourier constants could be obtained. Daily mean temperature values at both boundaries changed each day, causing slight discontinuity of boundary temperatures at midnight from day to day. This discontinuity had negligible effect on the numerical analysis.

The boundary conditions for water and solute are given in terms of net mass fluxes by

(7)

(8)

where E is the evaporation rate of water from the soil columns (m s^-1). For numerical analysis, the diurnal variation of E was calculated by spreading, via a sine function, the measured daily cumulative evaporation values across the daylight period. Zero evaporation rate from the soil columns was assumed for the night times.

Numerical Analysis
A numerical model developed by Nassar et al. (1997a) was used to solve the governing equations of heat, water, and solute (Eq. [1], [2], and [3]). An implicit finite difference scheme with a backward difference formula for the time derivative was used in the model. The model, along with the initial values of variables described by Eq. [4], the boundary conditions described by Eq. [5] through [8], and the transport parameters, was used to predict soil water content, temperature, and solute concentrations. Details are given in Nassar and Horton (1997) and Nassar et al. (1997a) for obtaining the transport parameters. The predicted values of water content and solute concentration were compared with the observed values.

Model and Results Evaluation
The ability of the model to predict observed values for water content and solute concentration was evaluated by computing model efficiency (Flerchinger et al., 1996). The model efficiency, which is the fraction of percentage of variation in measured values explained by the model, is computed as:

(9)

where eff is the model efficiency, V_o and V_p are the observed and predicted values, respectively, for water or solute content, V_m is the average of observed values of water or solute contents, and N is the number of observations. The significance among the treatments was evaluated by computing the F values (Mead et al., 1993). This test was used for comparing the cumulative evaporation from salinized and solute-free soils, and from compacted and noncompacted soils.

	Results and discussion

TOP NOTES ABSTRACT INTRODUCTION Materials and methods Results and discussion Conclusion REFERENCES

 
The observed values of cumulative evaporation, water content and solute concentration of the study are reported here as an average of three replications. The observed and predicted results are presented and discussed.

Figure 2 shows measured cumulative evaporation from salinized and solute-free soil columns. The cumulative evaporation was greater from solute-free soil than from salinized soil. The calculated F values (Mead et al., 1993) for the Clarinda and Fayette soils were 32.23** (significant at the 0.01 probability level) and 54.99**, respectively. The ratio of the cumulative water evaporations for salinized soil to solute-free soil increased with time from 0.78 to 0.89 for Clarinda soil and 0.9 to 0.95 for Fayette soil. The lowest evaporation ratios occurred during the first 24 h when the soils were wettest. The relatively large ratio for Fayette may be due to the relatively small initial solute concentration (0.92 mol kg^-1 of soil solution), water content (0.181 m³ m^-3), and specific surface area (120 m² g^-1) compared with the salinized Clarinda soil. Differences between Fayette and Clarinda soil in surface cracking due to drying may have affected cumulative evaporation also. Experimental and theoretical reductions of water transfer in salinized soil have been reported by Globus (1983), Nassar and Horton (1989), Bear and Gilman (1995), and Shimojima et al. (1996). Presence of solute in soil solutions reduces the equilibrium water vapor pressure (soil water potential) and may affect the water transport coefficients of soil (Nassar and Horton, 1997). Therefore, a salinized soil is expected to have less evaporation than solute-free soil. Salhotra et al. (1985) found that the reduction of evaporation from salinized water depended on saturation vapor pressure and the ionic composition of solute. The calculated vapor transport coefficients of Fayette soil decreased as solute concentration increased (Nassar and Horton, 1997). Similarly, Globus (1983) reported a reduction in the thermal water diffusivity of Podzolic salinized soil (Russian soil) compared with Podzolic solute-free soil. As the duration of evaporation increased, the soil dried and the water evaporation rate decreased. This reduction in the evaporation rate might be due to reductions of water vapor pressure and soil water transport coefficients that are in turn caused by reduction of water potential (soil water matric potential and solution osmotic potential). Predicted distributions of the soil water matric and solution osmotic potentials for noncompacted, salinized Clarinda and Fayette soils are shown in Fig. 3 . According to these distributions, the vapor pressure and soil water transport coefficients in the upper portion (0–0.05 m) were more affected by solution osmotic potential than by soil water matric potential after 2 d of evaporation. On the other hand, the soil water matric potential effects on water evaporation increased as time of evaporation increased. In the soil portion of 0.1 to 0.20 m, solution osmotic potential effects on the vapor and water transport coefficient of water were greater than the effects of soil water matric potential at 28 d of evaporation.

View larger version (24K): [in this window] [in a new window]   Fig. 2 Measured cumulative evaporation from noncompacted salinized and solute-free soils, and the ratios of cumulative evaporation of salinized to cumulative evaporation from solute-free soil

View larger version (24K): [in this window] [in a new window]   Fig. 3 Calculated soil water matric (m) and solution osmotic (o) potentials distributions after 2 and 28 d of evaporation for noncompacted salinized soil columns

View larger version (21K): [in this window] [in a new window]   Fig. 4 Measured cumulative water evaporation from salinized compacted and noncompacted soil, and the ratio of cumulative water evaporation from noncompacted to cumulative water evaporation from compacted soil

View larger version (30K): [in this window] [in a new window]   Fig. 5 Measured and predicted soil water contents, , for noncompacted salinized soil columns and noncompacted solute-free soil columns

View larger version (18K): [in this window] [in a new window]   Fig. 6 Measured, predicted, and initial soil water contents, , for compacted salinized soil columns

View larger version (25K): [in this window] [in a new window]   Fig. 7 Measured and predicted solute concentration in soil solution for noncompacted salinized soil columns

View larger version (29K): [in this window] [in a new window]   Fig. 8 Measured, predicted, and initial solute concentrations in soil mass for noncompacted salinized soil columns

View larger version (27K): [in this window] [in a new window]   Fig. 9 (a) Measured and predicted solute concentration in soil solution for compacted salinized soil columns, and (b) measured, predicted, and initial solute concentrations in soil mass for compacted salinized soil columns

	Conclusion

TOP NOTES ABSTRACT INTRODUCTION Materials and methods Results and discussion Conclusion REFERENCES

	NOTES

TOP NOTES ABSTRACT INTRODUCTION Materials and methods Results and discussion Conclusion REFERENCES

 
Journal Paper no. 17149 of the Iowa Agric. and Home Econ. Exp. Stn., Ames; Projects no. 3262 and 3287.

, Significant at the 0.01 probability level.

Received for publication December 16, 1996.

	REFERENCES

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• Bendz D., Bengtsson L. Evaporation from an active, uncovered landfill. J. Hydrol. 1996;182:143-155.
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• Nassar I.N., Shafey H.M., Horton R. Heat, water, and solute transfer in unsaturated soil: II—Compacted soil beneath plastic cover. Transport in Porous Media. 1997;27:39-55 a.
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