Soil Water Dynamics

The capacity of soil to regulate the terrestrial freshwater supply is a fundamental ecosystem service. Water percolating through soil is filtered, stored for plant utilization, and redistributed across flow paths to groundwater and surface water bodies. As such, the sustainability of water resources (considering both quantity and quality) is directly influenced by soil. Thus, most aspects of terrestrial- and freshwater aquatic-life depend on hydrologic processes in soil (O'Geen et al. 2010).

Water dynamics in soil are governed by many factors that change vertically with depth, laterally across landforms and temporally in response to climate (Swarowsky, et al., 2011). This article focuses on soil water dynamics and introduces concepts of soil moisture storage, water flow and the soil properties that influence these processes. The learning objectives discussed here include: 1) an introduction to the soil water balance equation; 2) the factors and soil properties that govern water potential and plant available water holding capacity; and 3) soil morphologic features and classification methods to describe the fate of water in soil. These concepts are integrated with a case study that describes the interplay between hydrologic processes, water dynamics in soil, and soil genesis.

Stored water in soil is a dynamic property that changes spatially in response to climate, topography and soil properties, and temporally as a result of differences between utilization and redistribution via subsurface flow (Figure 1; Western et al. 1999). Changes in soil moisture storage can be generalized with a mass balance equation (eq. 1) as a result of the difference between the amount of water added and that which is lost (Hillel 1982).

Change in soil moisture storage = inputs - outputs (eq. 1)

Water content increases (positive change in storage) when inputs including precipitation or irrigation exceed outputs. Water content decreases (negative change in storage) when outputs such as deep percolation, surface runoff, subsurface lateral flow, and evapotranspiration (ET) exceed inputs (eq. 1).

Water storage and redistribution are a function of soil pore space and pore-size distribution, which are governed by texture and structure (Childs 1940). Generally speaking, clay-rich soils have the largest pore space, hence the greatest total water holding capacity. However, total water holding capacity does not describe how much water is available to plants, or how freely water drains in soil. These processes are governed by potential energy. Water is stored and redistributed within soil in response to differences in potential energy. A potential energy gradient dictates soil moisture redistribution and losses, where water moves from areas of high- to low-potential energy (Hillel 1982). When at or near saturation, soils typically display water potentials near 0 MPa. Negative water potentials arise as soil dries resulting in suction or tension on water allowing the soil to retain water like a sponge.

Three soil moisture states, saturation, field capacity and permanent wilting point are used to describe water content across different water potentials in soil and are related to the energy required to move water (or extract water from soil) (Figure 2). When the soil is at or near saturation the direction of the potential energy gradient is downward through the soil profile or laterally down slope. This mechanism of flow by the force of gravity occurs mainly in macropores. As the soil dries, field capacity is reached after free drainage of macropores has occurred. Field capacity represents the soil water content retained against the force of gravity by matric forces (in micropores and mesopores) at tension of -0.033 MPa (Figure 2). As water content decreases, soil matric potential decreases, becoming more negative, and as a result, water is held more strongly to mineral surfaces due to cohesive forces between water molecules and adhesive forces associated with water and mineral particles (capillary forces). Water held between saturation and field capacity is transitory, subject to free drainage over short time periods, hence is it is generally considered unavailable to plants. This free water is termed drainable porosity. In contrast, much of the water held at field capacity is available for plant uptake and use through evapotranspiration (Figure 2).

The point at which matric forces hold water too tightly for plant extraction (-1.5 MPa) is termed the permanent wilting point. The amount of water held between field capacity and permanent wilting point is considered plant available water (PAW) (Figure 2). Water held between these two states is retained against the force of gravity, but not so tightly that it cannot be extracted by plants. Mesopores and micropores supply most plant available water. Water held at potentials below permanent wilting point (< -1.5 MPa) is not available for use by most plants because it strongly adheres to mineral particles. Water held at permanent wilting point is associated with partially filled micropores and hydrated surfaces of soil particles (Figure 2).

Differences in soil properties (texture and structure) affect the water content at saturation, field capacity, and permanent wilting point. Texture and structure determine pore size distribution in soil, and therefore, the amount of PAW. Figure 3 illustrates how the magnitude of PAW changes with soil texture. Coarse textured soils (sands and loamy sands) have low PAW because the pore size distribution consists mainly of large pores with limited ability to retain water. Although fine textured soils have the highest total water storage capacity due to large porosity values, a significant fraction of water is held too strongly (strong matric forces/low, negative water potentials) for plant uptake. Fine textured soils (clays, sandy clays and silty clays) have moderate PAW because their pore size distribution consists mainly of micropores. Loamy textured soils (loams, sandy loams, silt loams, silts, clay loams, sandy clay loams and silty clay loams) have the highest PAW, because these textural classes give rise to a wide range in pore size distribution that results in an ideal combination of meso- and micro-porosity. Soil structure can increase PAW by increasing porosity. Soil depth and rock fragment content also affect water holding capacity because bedrock and rock fragments are assumed to be unable to hold plant available water and/or accommodate plant roots.

Water movement in soil is closely linked with storage because water potential is a function of water content (Richards 1931). Water flow is also influenced by texture and structure, and other factors such as the layering of soil profiles. The rate of water flow is a function of the potential energy gradient and the ease with which water is transmitted through soil, termed saturated hydraulic conductivity, which is governed by pore size distribution and tortuosity of flowpaths (Childs & Collis-George 1950, Vogel 2008). Clay-rich soils have low saturated hydraulic conductivity due to a highly tortuous flow path. Conversely, sandy soils have larger pores and lower tortuosity that facilitate rapid water flow. The National Cooperative Soil Survey identifies the following permeability classes based on soil texture (Table 1). These classes can be modified (qualitatively) by the degree of soil structure. Strong soil structure, consisting of very fine and fine aggregates (e.g., granular, fine and medium angular blocky and subangular blocky) facilitate rapid drainage of soil by increasing macroporosity. In contrast, weak structure or coarse-sized structural units (prismatic or blocky) and platy structure can inhibit flow, creating a more tortuous flow path constraining water to inter-structural voids (Figure 4). Soil structure is highly relevant to water management in soils because it is subject to change either through deterioration by improper management, or to improvement through additions of soil organic matter. In contrast, it is usually infeasible to change texture.

Permeability Class	Permeability (cm/hr)	Textural class
Very slow	<0.13	clay
Slow	0.13–0.5	sandy clay, silty clay
Moderately slow	0.5–2.0	clay loam, sandy clay loam, silty clay loam
Moderate	2.0–6.3	very fine sandy loam, loam, silt loam, silty clay loam, silt
Moderately rapid	6.3–12.7	sandy loam, fine sandy loam
Rapid	12.7–25.4	sand, loamy sand
Very Rapid	>25.4	coarse sand
Table 1: Soil permeability classes and estimates of permeability rates by textural class.

The ease with which water drains from soil is equally as important as storage. For example, most terrestrial plants need to assimilate oxygen through roots, but oxygen is scarce in saturated soils. Moreover, microbial decomposition of organic matter is greatest (by orders of magnitude) under aerobic conditions. Poorly drained soils have limitations for a variety of land use practices. The recognition of poor drainage in soils is also used in wetland delineation efforts.

Drainage capacity can be identified through careful observation of soil properties. Poorly drained soils result in episodes of prolonged saturation, while excessively drained soils commonly experience water deficits. Soils that are saturated at times when the system is above biological zero (> 5° C) often develop redoximorphic features (Vepraskas & Sprecher 1997). The most common redoximorphic features in soil are iron and manganese concentrations and iron depletions (Figures 5 and 6). These features arise through microbial decomposition of soil organic matter under anaerobic conditions. Anaerobic conditions arise because the diffusion of oxygen in saturated soil is very slow and does not keep up with oxygen demands of aerobic respiration by microbes. When oxygen is depleted, facultative microbes utilize iron (Fe ³⁺) and manganese (Mn⁴⁺) as terminal electron acceptors to make energy. In doing so, these elements are reduced and become soluble in soil solution. As water mobilizes these soluble constituents they eventually encounter air (e.g., in root channels or other macropores) where they oxidize and re-precipitate as iron and manganese concentrations (Figures 5 & 6). Iron concentrations are usually rust colored, red or orange. Manganese concentrations are gunmetal blue, almost black (Figure 6). Redox depletions arise in microsites of extreme reducing conditions where Fe³⁺ has been completely removed. Redox depletions are typically dull in color (low Munsell chroma usually 2 or less for taxonomic purposes) and have blue, green, pale brown or yellow hues that reflect the coloration of primary minerals (Figure 7). The depth at which redoximorphic features occur is used to describe the extent of saturated conditions within a soil profile (Jacobs et al. 2002).

The National Cooperative Soil Survey describes five soil drainage classes determined by the rate at which water drains from soil and the height of the water table during the growing season (Table 2). Drainage classes are used for a variety of land-use decisions such as soil suitability for groundwater banking, land application of wastes, engineering and construction, septic system development, crop selection, land capability classification and wetland habitat.

Drainage Class	Water Removed	Redoximorphic features	Water table Height (m)
Excessively drained	Very rapid	None	>1.5
Somewhat excessively drained	Rapid	None	>1.5
Well drained	Readily not rapidly	Below 1 m	1.0–1.5
Moderately well drained	Somewhat slowly	Below 0.5 m	0.5–1.0
Somewhat poorly drained	Slowly	Below 0.25 m	0.25–0.5
Poorly drained	Very slowly	Gleyed or at or near soil surface	<0.25
Very poorly drained	Persistent saturation throughout the growing season
Table 2: Diagnostic features and conditions of soil drainage classes.

The amount and timing of precipitation ultimately governs soil moisture content, availability and flow. The temporal nature of moisture dynamics dictates ecosystem response and land-use decisions. Soil moisture regimes are used in Soil Taxonomy to describe annual variability in moisture as dictated by climatic factors and indirectly by soil and landscape factors. There are five general soil moisture regimes: aquic, udic, xeric, ustic, and aridic. These moisture regimes have detailed definitions (Soil Survey Staff 1999). General working definitions are described below.

Aquic - poorly drained soils that are saturated when the soil temperature (at 50 cm) is above biological zero (>5°C). These soils typically display evidence of prolonged saturation in the form of redoximorphic features in the root zone.

Udic - soils typical of humid or tropical environments that receive an even distribution of precipitation throughout the year. Soils are never dry for as long as 90 cumulative days or 60 consecutive days following summer solstice.

Ustic - soil moisture conditions intermediate between aridic and udic and typical of semiarid, tropical and monsoon climates, where soil is moist during part or all of the growing season followed by a prolonged dry season at some point in the year.

Xeric - a soil moisture regime with wet winters and dry summers typical of Mediterranean type climates.

Aridic - a soil moisture regime that is dry (low plant available water) for most times of the year.

Figure 8 depicts these soil moisture regimes across the globe. Soil moisture regimes serve as a broad scale planning tool to inform a variety land use decisions related to the hydrologic cycle such as, regions suitable for groundwater banking, dryland agriculture, summer fallow, or requiring irrigation technology, engineering considerations, and drainage infrastructure. These decisions cannot be made from climatic conditions alone. Soil moisture regimes are more useful in this type of decision making because it accounts for how soil properties (such as PAW) affect moisture dynamics within the soil profile.

The type of soil moisture regime can be used to infer the direction of water movement in soil, which affects soil development and morphology. The following case study illustrates this point. Figure 9 shows two soils, Mollisols (Xerolls and Ustolls) with 450 mm of mean annual precipitation that differ in the timing of precipitation (i.e., the soil moisture regime). Mollisols are typically described as grassland soils that have large stocks of below ground soil organic matter. The Ustoll is a Mollisol with an ustic soil moisture regime that formed in a bunchgrass prairie east of the Rocky Mountains (Figure 9a). The Xeroll is a Mollisol with a xeric soil moisture regime that formed in a bunchgrass prairie on the Columbia Plateau in eastern Washington (Figure 9b). These soils share similarities in soil forming factors including parent material, vegetation, age, and topography. The notable difference between these two profiles is the nature of the B horizons (subsurface zones of accumulation of clay, salts or iron). The Ustoll has a Bk horizon, which is a zone of accumulation of calcium carbonate (CaCO₃) (Figure 9a). This semi-soluble salt tends to accumulate in subsoils where deep percolation is minimal (Jenny 1994). The Xeroll has a Bw horizon, a subsoil where salts have been removed by deep percolation and the development of soil structure and slight accumulation of clay are detectable.

Potential evapotranspiration and precipitation over time in these soils illustrate the effects of timing of rainfall on soil moisture dynamics (Figure 10). In the Xeroll, precipitation is low during the growing season facilitating a period of utilization of stored water during the early growing season followed by a prolonged deficit (Figure 10a). Soil moisture is recharged at the onset of fall rains when the dormant season begins. As precipitation continues throughout the winter months a surplus is achieved. Surplus occurs when soils become saturated and water is allowed to freely drain with the force of gravity. A surplus in soil moisture results in the loss of free water stored between saturation and field capacity, which is subject to gravitational flow. In this scenario there are two main flowpaths: 1) deep percolation of free water beyond the root zone, and 2) losses of stored water by ET during the growing season. The large proportion of surplus water present during the winter months when plants are dormant facilitates the leaching of soluble and semi-soluble salts out of the soil profile and into the groundwater.

In the Ustoll, most precipitation occurs during the growing season and is utilized through transpiration until a soil moisture deficit occurs (Figure 10b). As evapotranspiration decreases during the dormant season, soil recharge occurs, but precipitation is not high enough at this time to create a surplus. As a result, soil moisture is stored over the dormant season (winter months) and the major flow path is upward via capillary rise as plants become active in spring, extracting water by evapotranspiration. The absence of a surplus results in incomplete leaching of the soil profile and the accumulation of semi soluble salts (white coatings), as symbolized by the Bk horizon (Figure 9a).

This comparison of ustic and xeric soil moisture regimes provides an example of how soil moisture dynamics and timing of precipitation govern water utilization and movement in soil. Timing of precipitation is equally as important as its magnitude when considering soil water dynamics. This case study demonstrates that the fate of water in soil (e.g., deep percolation or evapotranspiration) governs the expression of soil morphologic features. Thus, soil features can be used to infer hydrologic processes.

Water storage dynamics and flow facilitate the four basic soil forming processes: translocations, transformations, additions and losses of soil constituents in a soil profile. These processes determine the chemical, morphological and physical properties of soil such as the variation of texture with depth. Hydrological processes active in soil contribute to weathering processes, and indicators of these processes are preserved by the soil profile in the form of observable and measurable soil characteristics, similar to those discussed in the case study (O'Geen et al. 2002). Other soil morphologic indicators of hydrologic process include redoximorphic features, abrupt accumulation of clay in the subsoil, development of soil structure, and presence of cemented layers (Fritsch & Fitzpatrick 1994). Thus soil resource inventories (e.g., soil survey), which document soil properties, can be used to infer hydrologic processes.

Although climatic factors ultimately control plant-water relationships, soils regulate water acting as a sponge to hold water against gravitational forces in plant available form. Soil properties such as texture and structure govern pore size distribution, which dictates total water storage, available water holding capacity, and water movement in soil. While it is generally not feasible to modify soil texture to improve plant-water relationships, soil structure can be enhanced by adding organic matter to promote more meso- and macro-porosity, which increases plant available water holding capacity and helps to promote free drainage. An understanding of soil water relationships is fundamental to most land use decisions.