Soil Sampling and Methods of Analysis - Part 7 pot

VII. SOIL WATER ANALYSES Section Editors: W.D. Reynolds and G. Clarke Topp ß 2006 by Taylor & Francis Group, LLC. ß 2006 by Taylor & Francis Group, LLC. Chapter 69 Soil Water Analyses: Principles and Parameters W.D . Reynolds Agriculture and Agri-Food Canada Harrow, Ontario, Canada G. Clarke Topp Agriculture and Agri-Food Canada Ottawa, Ontario, Canada 69.1 INTRODUC TION For the uninit iated, soil water anal yses can be daunt ing becau se they are b ased on many nonin tuitive princi ples and they use a large numb er of comple x parame ters. The prima ry intent of thi s chapter is to help alleviate this situat ion by b riefly revi ewing the main principle s and para meters involved in mode rn soil water anal yses. The chapter also serves as addi tional backgr ound and context for the methods describ ed in Chapte r 70 thr ough Chapter 85. Soil water analyses can be organized into two main groups: (i) analysis of storage properties and (ii) analysis of hydraulic properties. Storage properties refer to the soil’s ability to absorb and hold water, and these properties include water content, water potential, and water desorption and imbibition characteristics. Hydraulic properties, on the other hand, refer to the soil’s ability to transmit or conduct water, and these include saturated hydraulic conduct- ivity, unsaturated hydraulic conductivity, and various associated capillarity parameters such as sorptivity, flux potential, sorptive number, and flow-weighted mean (FWM) pore diam- eter. These properties and their interrelationships are discussed in the following sections. 69.2 SOIL WATER CONTENT Soil water content can be defined on a gravimetric basis (mass of water per unit mass of dry soil) or on a volumetric basis (volume of water per unit bulk volume of dry soil), and it is expressed either as a dimensionless ratio or as a percentage. These two definitions are not equivalent, however, and it is consequent ly essential to specify the definition used when reporting water content values. It should also be noted that ‘‘bulk volume’’ of dry soil refers to the dimensions of the soil sample just before the water volume determination and before ß 2006 by Taylor & Francis Group, LLC. any soil disturbance . Gravi metric water cont ent (i.e., mass water =mass dry soil) is related to volumetric water cont ent (i. e., volume water =bulk volume dry soil) via soil dry bulk density , r b (mg m À 3 ) and pore water density , r w (mg m À 3 ), accordi ng to the formul a u v ¼ r b r w  u m (69: 1) where u m is the gravim etric water content (kg water kg soil À 1 ) and u v is the volumetric water content (m water 3 m soil À 3 ). The volumetric water content is often expre ssed as ‘‘relative saturati on’’ (al so know n as saturatio n ratio or degree o f saturati on), which gives the ratio of the measur ed volumet ric water content ( u v ) to the correspo nding volumet ric water cont ent at full or com plete saturatio n (u s ). Con sequentl y, relative sat uration give s the fraction o f the soil pore space that is water -filled, and therefore ranges from a minimum valu e of 0 (no water in pore space) to a max imum value of 1 (pore space comple tely water -filled). When the soil pore space is completely water-fille d (relati ve saturati on ¼ 1), the soil volumet ric water cont ent is equal to the soil porosity, n (por osity is defined as the total volume of soil pore space per unit bulk volume of soil). Relativ e saturati on is frequently expre ssed as an ‘‘effect ive saturati on,’’ S e , which include s the residua l soil water content , u r , that is the ‘‘im mobile’’ water remainin g in air-dry soil and retained in small isolat ed pores. Effective sat uration is defi ned as S e ¼ ( u v Àu r ) =( u s Àu r ) and ranges from a minimum v alue o f 0 at residua l saturati on (i.e., u v ¼u r ) to a max imum valu e of 1 at complete saturation (i.e., u v ¼u s ). Water content measur ement tech niques are often classified as ‘‘direct ’’ or ‘‘indirect.’ ’ Direct methods usually alter the sample irrevocabl y by changing its water content and physi cal characterist ics (i.e., they are ‘‘de structive ’’ methods ); and these methods involv e som e form of remo val or separation of water from the soil matrix with a dir ect measur ement of the amount of water removed . Separation of the water from the soil matri x may be achieved by heating (wate r vapor ization), by water repl acement with a solvent (w ater absor ption), or by chemical reac tion (water d isassociati on). Th e amo unt of water removed is then determin ed by measuring the change in soil mass after heating, by collecting and condensing emitted water vapor, by chemical or physical analysis of the extracting solvent, or by quantitative measurement of chemical reaction products. The removal of water by heating is commonly referred to as the thermogravimetric technique (see Topp and Ferre ´ 2002 for details) and it is by far the most com mon of the dir ect methods (see Chapter 70). Indirec t methods measur e some physical or chemical property of soil that depends on its water content. These properties include the relative permittivity (dielectric constant), electrical conductivity, heat capacity, hydrogen content, and magnetic susceptibility. The indirect methods usually alter the sample minimally (or not at all) in that the water content and physical characteristics of the sample are not changed appreciably by the measurement (i.e., they are ‘‘nondestructive’’ methods). However, the accuracy and precision of indirect methods depends to a large extent on the accuracy and precision of the relationship between the measured property (e.g., permit- tivity) and u v . In Chapter 70, we limit discussion to the indirect methods that are based on relative bulk soil dielectric permittivity, as they are the most highly developed and versatile. The electromagnetic (EM) methods discussed in Chapter 70 all arise from analyses based in EM wave propagation or radio frequency (RF) circuits. Measurement of soil water content by these methods involves using the soil as an EM wave-propagating medium or as a resistor or capacitor in a circuit. The time-domain reflectometry (TDR), ground-penetrating radar (GPR), and remote radar (remote sensing) methods use the EM wave-propagation properties ß 2006 by Taylor & Francis Group, LLC. of the soil, whereas the capacitanc e and impeda nce methods use the soil as a resist or o r capac itor in a circuit. The unique electrical properties of water (both pure water and soil pore water) form the basis of soil water content measurements by EM wave propagation. The relative dielectric permittivity of water is generally more than an order of magnitude larger than that of other soil components. As a result, the bulk relative dielectric permittivity of soil ( « ra ) is almost entirely a function of the soil’s volumetric water content (u v ), with only a slight dependence on the volume fraction of soil solids and the bulk soil electrical conductivity (Topp et al. 1980). For each of the EM methods presented in Chapter 70, a measurement of « ra is used to infer u v . A single relationship between « ra and u v for all soils does not yet exist because of the complex interactions among EM waves and soil component s. Many soils have very similar relationships, however, and thus sufficient accuracy can usually be attained using only a few ‘‘quasigeneral’’ relationships, for example, mineral soils, organic soils, saline soils, etc. (Topp et al. 1980; To pp and Ferre ´ 2002). It has further been established that assuming a linear relationship between u v and ffiffiffiffiffiffi « ra p is appropriate for most soil materials (Topp and Reynolds 1998), and that this relationship is predicted by dielectric models employing a three-component mixture of soil solids, soil water, and soil air (Robinson et al. 2005). Thus EM methods are now considered highly reliable for measuring soil volumetric water content. Water content methods are described in Chapter 70 and include thermogravimetry (oven drying), TDR, GPR, and a general description of the impedance and capacitance techniques. Thermogravimetry based on oven drying is usually considered the ‘‘benchmark’’ of accuracy and relevance against which other methods are assessed. 69.3 SOIL WATER POTENTIAL Total water potential (c t ) is classically defined as the amount of work (force Âdistance) required to transport, isothermally and reversibly, an infinitesimal quantity of water from a specified reference condition (pool of pure water at specified pressure and elevation) to the system under consideration (Or and Wraith 2002). It is usually more convenient for natural porous materials, however, to consider c t as the amount of work required to transport water away from the material (i.e., remove water rather than add water), as most natural materials are hydrophilic and thereby tend to absorb and retain water in a manner similar to that of a paper towel (nonswelling materials) or sponge (swelling materials). Water potential is commonly expressed in units of energy per unit mass, U m (J kg À1 ), energy per unit volume, U v (Pa), or energy per unit weight, U wt (m), with the latter two being by far the most prevalent. Conversion amongst the units is achieved using U m ¼ U v =r w ¼ gU wt (69:2) where r w is the density of water (1000 kg m À3 at 20  C) and g is the acceleration due to gravity (9:81 m s À2 ). The total potential, c t , of water in soil or other natural porous materials is usually the sum of four-component water potentials: c t ¼ c m þ c p þ c p þ c g (69:3) ß 2006 by Taylor & Francis Group, LLC. where c m is the matric pote ntial, c p is the osm otic pote ntial, c p is the pressure pote ntial, and c g is the gravit ational pote ntial. Th e matri c potential is negativ e (c m 0) and arises from the various electrostat ic forc es in the soil matri x that attract water when the soil is u nsaturated . The osmoti c potential is als o negativ e ( c p 0) and resu lts from disso lved mater ials (salts) and colloids, which lower the pore water’s activit y (free energy) below that of pure water . The pres sure potential is posi tive ( c p ! 0), and is cause d by the hydros tatic pressure of the pore water overlying the measur ement point whe n the soil is sat urated. The gravitationa l potential ( c g ) arises from the action of the earth’ s g ravitational forc e field on the pore water and can be either positive or negative depending on whet her the datum is b el ow or above the measuring point, r espect ively. Campbe ll (1987) reviews the various techniques for measuring matric potential and t he type of sensors e mployed; and t he readers are recommended t o r efer to Passioura (1980) for a more detailed discussion of the m eaning of matri c potent ial. A reaso nable estim at e o f o smotic po tential c an be derived f rom measurements of electrical conductivi ty corrected f or water c ontent (Gupta and H anks 1972); howeve r, more reliable measures can b eobtainedbyextractingsoilporewater and m easuring the osmotic potential directly in a thermocouple psychrometer or by usi ng the combine d pr es sur e c h ambe r a nd thermocouple psychrometer system of Campbell (1987). W he n the matric, pressure, and gravitational p ot enti al s a re expressed in unit s o f e nergy p er unit w ei ght (U wt ), they are generally called ‘‘heads’’ rather than potentials, a nd they are e quival ent to t he vertical distance between the m ea surement p o i n t ( e . g . , piezometer intake, tensiometer cup, etc.) and either the free surface water level (for matric and pressure heads) or the selected reference elevation or datu m (for gravitational heads) (Figure 69.1). Water flow can be induced by gradients in all four water potentials, although a gradient in osmotic potential requires the presence of a membrane that is permeable to water but impermeable to selected solutes and colloids (Or and Wraith 2002). Methods for measur ing water pote ntial are describ ed in Chapter 71 and include the piezometer method, the tensiometer method, resistance block methods, and selected thermocouple psychrometer methods. (a) Datum (water table) Piezometric surface Piezometer riser pipe Manometer Soil surface Selectively permeable cup (to water, not air) Tensiometer point of measurement Well screen Piezometer point of measurement z = 0 (b) p m t g g (usually sea level) t FIGURE 69.1. The operating principles of a piezometer (a) and a tensiometer (b). The piezometer measures pressure potential (c p ), and the tensiometer measures matric potential (c m ). ß 2006 by Taylor & Francis Group, LLC. 69.4 SOIL WATER DESORPTION AND IMBIBITION Soil water desor ption and imbib ition curves char acterize the rel ationship betwee n soil volumet ric water cont ent, u v [L 3 L À 3 ] (Chapter 72 through Cha pter 74), and pore water matric head, c m [L] (Chapt er 71). The desorption curve (also know n as the water release characteristic, water retention curve, and soil moisture characteristic) describes the decrease in u v from saturation as c m decreases from zero, whereas the imbibition curve describes the increase in u v from dryness as c m increases from a large negative value (see Figure 69.2). The two curves generally have different shapes because of hysteretic effects (Hillel 1980); and when a partially drained soil is rewetted, or when a partially wetted soil is redrained, the relationship between u v and c m usually follows an intermediate and nonunique path between the desorption and imbibition curves (see Figure 69.2). For this reason, the desorption curve is often referred to as the ‘‘mai n drainage curve’’; the imbibition curve as the ‘‘main wetting curve’’; and the intermediate curves as ‘‘scanning curves’’ (see Figure 69.2). When the soil has a relatively uniform and narrow pore size distribution (e.g., structureless sandy soil), distinct ‘‘air-entry’’ and ‘‘water-entry’’ matric heads can occur on the desorption and imbibition curves, respectively (Figure 69.2). The air-entry head or value, c a [L], is the pore water matric head where the saturated soil (i.e., u v constant and maximum) suddenly starts to desaturate as a result of decreasing c m ; and the water-entry head or value, c w [L], is the pore water matric head where an unsaturated soil suddenly saturates as a result of increasing c m . Both c a and c w are negative, and typically, jc a j%2jc w j (Bouwer 1978). Also note that in Figure 69.2 that the saturated volumetric water content on the imbibition curve (i.e., u fs at c m ¼ 0) is less than the saturated volumetric water content on the desorption curve (i.e., u s at c m ¼ 0), which is a consequence of air entrapment in soil pores during the wetting process (Bouwer 1978). As implied above, soil water desorption Imbibition or main wetting curve Scanning curves Desorption or main drainage curve 0 q fs q s q +− a w FIGURE 69.2. Desorption, imbibition, and scanning curves, u(c), for a hysteretic soil. The arrows indicate the direction of the drainage and wetting processes. Note that the satur- ated volumetric water content for the imbibition curve, u fs , is less than that for the desorption curve, u s , due to air entrapment upon rewetting. Note also that the water-entry matric head, c w [L], is greater (less negative) than the air-entry matric head, c a [L]. ß 2006 by Taylor & Francis Group, LLC. and imbib ition is a com plicated proce ss that is difficul t and time-con suming to char acterize in d etail. Fort unately, it is usually not necessar y to measur e the scann ing curve s and Cha pter 72 through Chapter 74 conse quent ly focus on dete rmination of only the desor ption (main drainage) curve and the imbibition (main wetting) curve. 69.4.1 APPLICATION OF DESORPTION AND IMBIBITION CURVES The shape and magnitude of desorption and imbibition curves depends on the number and size distribution of the soil pores, which in turn depends on texture, porosity, structure, organic matter content, and clay mineralogy. Figure 69.3 gives schematic examples of desorption curves for a representa tive coarse-textured, unstructured soil (e.g., uniform sandy soil), and for a representative fine-textured soil (e.g., clayey soil) with and without structure, where ‘‘structure’’ refers to the presence of aggregates, peds, and macropores (i.e., large cracks, root channels, worm holes, etc.). Note that the coarse-textured (sandy) soil retains less water than the fine-textured (clayey) soil (i.e., lower u v values), and it also releases its water in a different manner (i.e., different curve shape). Note also that soil structure can increase the saturated water content (if the bulk density decreases) and it can cause the wet-end of the desorption curve to be very steep relative to a structureless condition when aggregates, peds, and macropores are not present. Soil water desorption and imbibition curves are important for determining soil pore size distribution, for interpreting soil strength data, and for determining the transmission and Structured clayey soil Unstructured clayey soil Unstructured sandy soil 0 + q s q s q s − FIGURE 69.3. Soil water desorption curves for a ‘‘representative’’ unstructured sandy soil, and a representative clayey soil with and without structure. u s [L 3 L À3 ] is the satur- ated volumetric water content and c [L] is pore water matric head. Note that the increase in u s for the structured clayey soil relative to the unstructured clayey soil implies a decrease in soil bulk density. If bulk density remains constant, the presence of structure changes only the shape of the curve and not the value of u s . ß 2006 by Taylor & Francis Group, LLC. storage of flu ids (liquids, gases ) in the soil profile. The sizes of soil pores relevan t to the storage and transmis sion of fluids are determin ed from desorption and imbib ition curves via the Kelvin or ‘‘capi llary rise’’ equat ion. Soil strengt h relat ionships , such as cone penet ration resistanc e and vane shear, are h ighly depend ent on the antecede nt soil water cont ent at the time o f the mea sureme nt, and must therefore be related to the desorption and imbib ition curve s befor e deta iled anal yses can be conduc ted. With resp ect to water and solu te transm ission, the desor ption and imbibition curve s are require d for defin ing the water capacity relationship in the water transport (Richards) equation, and various solute sorption– desorption relationships in the solute transport (convection–dispersion) equation. With respect to water and air storage , the desorption and imb ibition curves are used to determ ine saturated and field- saturated soil water content s, field capacity water cont ent, permanent wiltin g point water content , air capacity, and plant- available water capac ity. These water =air storage para meters and othe r quantitie s derived from these parameter s are defined and briefly discusse d in the follow ing sectio ns. 69.4.2 WATER AND AIR STORAGE PARAMETERS The volumetric water content, u v [L 3 L À3 ], for a rig id soil (i.e., no shrinkage or swelling) is defined by u v ¼ V w =V b (69:4) where V w [L 3 ] is the volume of soil water per unit bulk volume of dry soil, V b [L 3 ] (see Sect ion 6 9.2). When the soil is com pletely saturated (i.e., no entrappe d air), V w ¼ volume of pore space and thus u v ¼ u s ¼ soil porosity. When the soil is ‘‘field-saturated’’ (entrapped air present), V w < volume of pore space and u v ¼ u fs < soil porosity, usually by 2–5 percentage points (Bouwer 1978). For most field applications where wetting and drying are involved, u fs is a more relevant measure of the maximum soil volumetric water content than u s or porosity because entrapped air is almost always present. Field water capacity (more commonly known as field capacity, FC) is formally defined as the amount of water retained in an initially saturated or near-saturated soil after 2–3 days of free gravity drainage without evaporative loss (Hillel 1980; Townend et al. 2001). For application purposes, however, FC is usually defined as the equilibrium volumetric water content, u FC , at a specified matric head, c FC . For intact soil containing normal field structure, c FC ¼À1 m is most often used, although values as high as c FC ¼À0:5 m have been recommended for wet soils with a shallow water table, and as low as c FC ¼À5 m for dry soils with a very deep water table (Cassel and Nielsen 1986). If the soil has been disturbed and repacked, use of c FC ¼À3:3 m is usually considered to provide u FC values that are comparable to intact soil values. The permanent wilting point (PWP) is defined as the soil water content at which growing plants wilt and do not reco ver when the evapotranspirative demand is eliminated by providing a water vapor–saturated atmosphere for at least 12 h (Hillel 1980; Romano and Santini 2002). Once the soil water decreases to the PWP value, plants are permanently damaged and may even die if water is not added quickly. In this respect, the PWP water content also represents the amount of ‘‘plant-unavailable’’ water; i.e., water that is too strongly held by the soil to be extracted by plant roots. Although the true PWP can vary widely with plant species, plant growth sta ge, and soil type, it has been found that the equilibrium volumetric water content, u PWP , at the matric head, c PWP ¼À150 m, is a ß 2006 by Taylor & Francis Group, LLC. suitable working definition (Soil Science Society of America 1997). This is because water content becomes relatively insensitive to matric head (i.e., water content is nearly constant) in the c m À150 m range for most agricultural soils (Romano and Santini 2002). Plant growth and performance is critically dependent on adequate supplies of air and water in the root zone. Convenient and popular measures of the soil’s ability to store and provide air and water for plant use are the so-called air capacity and plant-available water capacity. Air capacity (AC) is defined as AC ¼ u s À u FC (69:5) and proposed minimum values for adequate root-zone aeration are 0:10 m 3 m À3 for loamy soils (Grable and Siemer 1968), 0 :15 m 3 m À3 for clayey soils (Cockroft and Olsson 1997), and about 0:20 m 3 m À3 for horticultural substrates (Verdonck et al. 1983; Bilderback et al. 2005). Field soils that have AC values appreciably below these minimums are susceptible to periodic and damaging root-zone aeration deficits. Plant-available water capacity (PAWC) is defined as PAWC ¼ FC À PWP (69:6) and it represents the maximum amount of water that a fully recharged soil can provide to plant roots. This definition is based on the concept that soil water at c m > c FC drains away too quickly to be captured by plant roots, whereas water at c m < c PWP is held too strongly by the soil to be extracted by the roots (compare PWP discussion). The propos ed minimum PAWC for optimum plant growth and minimum susceptibility to droughtiness is 0.20–0.30 m 3 m À3 (Verdonck et al. 1983; Cockroft and Olsson 1997; Bilderback et al. 2005). Recent research (Olness et al. 1998; Reynolds et al. 2002) suggests that the optimal balance between root-zone soil water and soil air is achieved in rain-fed crops when FC=Porosity ¼ 0:66 (69:7) or alternatively, when AC=Porosity ¼ 0:34 (69:8) These criteria are based on the finding that maximum production of crop-available nitrogen by aerobic microbial mineralization of organic matter occurs when about 66% of the soil pore space in the root zone is water-filled, or alternatively, when 34% of the pore space is air- filled (Skopp et al. 1990). The rationale for applying Equation 69.7 and Equation 69.8 to rain-fed crops is that root-zone soils with these ratios are likely to have desirable water and air contents (for good microbial production of nitrogen) more frequently and for longer periods of time (especially during the critical early growing season) than root-zone soils that have larger or smaller ratios. 69.4.3 DETERMINATION OF DESORPTION AND IMBIBITION CURVES The generally accepted ‘‘ideal’’ for obtaining soil water desorption and imbibition curves is to collect simultaneous field-based measurements of volumetric water content, u v , and ß 2006 by Taylor & Francis Group, LLC. [...]... N., and Durner, W 2002 Simultaneous determination of water transmission and retention properties: inverse methods In J.H Dane and G.C Topp, Eds., Methods of Soil Analysis, Part 4— Physical Methods, Soil Science Society of America, Madison, WI, pp 963–1004 Campbell, G.S 19 87 Soil water potential measurement In R.J Hanks and R.W Brown, Eds., Proceedings of International Conference on Measurement of Soil. .. Klute, Ed., Methods of Soil Analysis, Part 1—Physical and Mineralogical Methods 2nd ed American Society of Agronomy, Madison, WI, pp 901–926 Koorevaar, P., Menelik, G., and Dirksen, C 1983 Elements of Soil Physics Elsevier, New York, 228 pp Cockroft, B and Olsson, K.A 19 97 Case study of soil quality in south-eastern Austrialia: ß 2006 by Taylor & Francis Group, LLC McIntyre, D.S 1 974 Soil sampling techniques... 15: 74 7 75 1 Bouma, J 1983 Use of soil survey data to select measurement techniques for hydraulic conductivity Agric Water Manage 6: 177 –190 Bouma, J 1985 Soil variability and soil survey In J Bouma and D.R Nielsen, Eds Proceedings of Soil Spatial Variability Workshop PUDOC, Wageningen, The Netherlands, pp 130–149 Bouwer, H 1966 Rapid field measurement of air-entry value and hydraulic conductivity of soil. .. flow system analysis Water Resour Res 2: 72 9 73 8 Bouwer, H 1 978 Groundwater Hydrology McGraw-Hill, Toronto, ON, Canada Bruce, R.R and Luxmoore, R.J 1986 Water retention: Field methods In A Klute, Ed., Methods of Soil Analysis, Part I—Physical and Mineralogical Methods 2nd ed American Society of Agronomy, Madison, WI, pp 663–683 management of structure for roots in duplex soils In E.G Gregorich and M.R... Sci 85: 228–232 Grable, A.R and Siemer, E.G 1968 Effects of bulk density, aggregate size, and soil water suction on oxygen diffusion, redox potentials, and elongation of corn roots Soil Sci Soc Am Proc 32: 180–186 Gupta, S.C and Hanks, R.J 1 972 Influence of water content of electrical conductivity of the soil Soil Sci Soc Am Proc 36: 855–8 57 Hillel, D 1980 Applications of Soil Physics Academic Press,... infiltration and seepage In Y.S Fok, Ed., Infiltration, Development and Application Water Resources Research Centre, Honolulu, HI, pp 1– 27 Romano, N and Santini, A 2002 3.3 Water Retention and Storage, 3.3.3 Field In J.H Dane and G.C Topp, Eds., Methods of Soil Analysis, Part 4—Physical Methods Soil Science Society of America, Madison, WI, pp 72 1 73 8 Raats, P.A.C 1 976 Analytical solutions of a simplified... Resour Res 16: 574 –582 ´ Topp, G.C and Ferre, Ty P.A 2002 3.1 Water content In J.H Dane and G.C Topp, Eds., Methods of Soil Analysis, Part 4—Physical Methods Soil Science Society of America, Madison, WI, pp 4 17 545 Topp, G.C and Reynolds, W.D 1998 Time domain reflectometry: a seminal technique for measuring mass and energy in soil Soil Till Res 47: 125–132 Townend, J., Reeve, M.J., and Carter, A 2001... Permittivity Of interest for water content determination is the two-way travel time of the TDR signal in the soil in and surrounding the probe Two times are measured; the time of arrival of signal reflected from the probe-to -soil interface (t1 in Figure 70 .2) and the time of arrival of the signal reflected from the end of the probe (t2 in Figure 70 .2) The TDR waveform in Figure 70 .2 shows the recommended way of. .. (2002), Reynolds and Elrick (2005), and references contained therein Saturated hydraulic property methods are described in Chapter 75 through Chapter 79 and Chapter 84; and they include the constant and falling head core methods (Chapter 75 ), selected constant and falling head well permeameter methods (Chapter 76 ), selected constant and falling head ring infiltrometer methods (Chapter 77 ), the auger hole... size and interconnectedness of water-conducting pores Alternatively, formation of precipitates and swelling=dispersion of silt and clay particles will usually decrease Ksat through narrowing and plugging of pores Reduction in Ksat most commonly occurs in silt- and clay-rich soils ß 2006 by Taylor & Francis Group, LLC when the cationic speciation is changed or the concentration of the resident soil . 79 and Chapter 84; and they include the const ant and falling head core methods (Chapter 75 ), selected const ant and falling h ead well permeame ter methods (Chapter 76 ), sel ected constant and. curves for a representa tive coarse-textured, unstructured soil (e.g., uniform sandy soil) , and for a representative fine-textured soil (e.g., clayey soil) with and without structure, where ‘‘structure’’. of structure changes only the shape of the curve and not the value of u s . ß 2006 by Taylor & Francis Group, LLC. storage of flu ids (liquids, gases ) in the soil profile. The sizes of soil