Chapter four Soil system

CHAPTER

FOUR

SOIL S Y S T E M

4.1

INTRODUCTION

As discussed in Chapter 1, the soil system, referred to as the geosphere, is an integral part of the land environment. Composed of soil constituents, water and air, it is the living environment for and from which human, animals and plants extract most of their food and energy. In this respect, the soil system is probably the ecosystem that receives the highest degree of anthropogenic stresses. From the geoenvironmental engineering viewpoint, the primary concerns relate to the problem of land disposal of waste and the decontamination of polluted soils. Evaluation of the effectiveness of a good barrier system, for land disposal, can usefully benefit from a closer consideration of how the pollutants are partitioned within the soil system. The pollutant retention mechanisms vary with soil constituents (mineral, amorphous material, soil organic matter, carbonates, etc.). A precise knowledge of a soil's retention mechanism can be used to estimate: (1) The potential for geoaccumulation, i.e., the tendency of a pollutant to persist in the soil for a long time; (2) The potential bioavailability of pollutants; bioavailability refers to the fraction of the total pollutant that is available for uptake, from polluted soils, by biota; and (3) The appropriate decontamination process, i.e., physical, chemical, biological, or electrical. The following development of the soil system will be concerned with attributes and characteristics which are pertinent to its utility as a waste retention and/or decontaminating agent.

4.2

SOIL PHASES

4.2.1

Gas Phase

The gas phase in soils is called soil air. It is located in the pore spaces, defined as the space in soil that is occupied by the gas and liquid phases. Soil air is composed of the same type of gases commonly found in the atmosphere. Biological activity in soil, however, may cause the percentage composition of soil air to differ considerably from that of atmospheric air [781 ml nitrogen (N2), 209 ml oxygen (O2) , 9.3 ml argon (Ar), and 0.31 ml carbon dioxide (CO2) in one litre of dry air]. Wellaerated soil contains 180-205 ml O2 per liter of soil air, but may drop to 100 ml/1 at one metre below the soil surface after a rainfall (i.e., flooding of the soil). The fractional volume of CO2 in soil air is typically 3-30 ml/1, but can approach 100 ml/1 at one metre depth after the flooding of the soil. The high CO2 content of soil air, relative to that of the atmosphere, can have a significant impact on a soil's acid buffering capacity, i.e., the ability of the soil to resist changes in its pH. 59

60

SOIL SYSTEM

The dissolution of soil air gases into the soil solution is an important factor in the cycling of chemical elements in the soil environment and the remediation of polluted soil by vapour extraction. Gaseous species are partitioned between soil air and soil water. At equilibrium, the relation between the concentration of gas in soil solution and the partial vapour pressure in soil air is given by Henry's law:

[4.1]

KH = [Aaq]/eA

where K/4 is Henry's constant (mol/m3), [Aaq] is the concentration of gas A in the soil solution (mol/m3), and PA is the partial vapor pressure of gas A in the soil air (atm). Eq. [4.1 ] is only valid when the gas concentration in the soil solution is low. The importance of Eq. [4.1] lies in its use as a method for: (1) classifying the potential volatilization of organic chemicals, (2) modelling the partitioning of organic chemicals in unsaturated soils, and (3) defining the limits for using soil vapour extraction process to decontaminate soils polluted with organic chemicals. Table 4.1 lists values of KHat 25~ for several uncontaminated soil air gases. For example, from Table 4.1, K, = 34.06 mol/m 3 at 25~ for CO). Assuming the partial pressure of CO2 in soil air is 0.03 atm; then, according to Eq. [4.1 ], the concentration of [CO2 (aq)] in the soil solution is 1.02 mol/m 3.

Table 4.1: Values of Henry's constant of various ~ases in soil at 25~

x.

K. Gas CO 2 CH 4 NH 3

N20

mol

m 3 atm 1

Gas

34.06 1.50 5.76• 104 25.55

NO 02 SOz HzS

mol

m -3

atm ~

1.88 1.26 1.24• 103 1.02•

Aeration has a significant effect on many biochemical reactions in soils. Many of the decomposition reactions of soil organic matter are influenced by the soil's aeration status. When adequate amount of air (02) is present, aerobic reactions (oxidation) prevail; when aeration is poor, anaerobic reactions (reduction) predominate. It is noteworthy that soil properties in oxidized and reduced states are markedly different. Thus, the solubility of pollutants in soils is dependent on the oxidation-reduction state, and has an important impact on the mobility of pollutants through soils. In well-aerated soils, organic matter will decompose into CO2 and H20, and release nutrient elements. The carbon dioxide produced will react with soil water to form carbonic acid, and alter the soil acidity. The increase in soil acidity will enhance the dissolution of soil minerals. Other acids that are produced by aerobic decomposition of organic matter include humic, nitric, and sulfuric acids. These acids also contribute to the solubilization of soil minerals. Such dissolution is disadvantageous

SOIL PHASES

61

for the design of land disposal systems since it increases soil hydraulic conductivity and reduces the ability of soils to retain pollutants. Reduction processes, which prevail in anaerobic conditions, contribute to the reduction of many soil pollutants, e.g., reduction of iron to Fe 2+, manganese to Mn 2+, sulphur to SO32-, and nitrate to nitrite (NO2). Through a series of successive reduction, a nitrogen compound may eventually be reduced to N2 gas. This process, known as denitrification, is often applied to remove nitrates from polluted soil, as discussed in Chapter 21. Soil aeration is affected by compaction, soil moisture and temperature. When the soil pores are destroyed by compaction, not only will air be deficient, but also, the soil becomes dense. As the soil moisture content decreases, the soil air content increases, reaching its maximum (15% 02) at the field capacity. Field capacity refers to soil moisture content at a moisture tension of 0.3 bars (Brady, 1984). With increased soil drying, the air content remains constant. As the temperature increases, the 02 concentration is increased. 4.2.2

Fluid Phase

The fluid phase in soil constitutes between one- and two-thirds of the total soil volume. Soil water exists principally in the condensed phase. Soil water is a repository for dissolved solids and gases and, for this reason, is commonly referred to as soil solution. The term soil moisture is frequently used to refer to soil water or soil solution. Soil moisture can be found in the macro- and micro-pores. When water saturated soil drains under the influence of gravity, some water is retained in the macro-pores in the form of a thin film around the soil particles. The micro-pores, on the other hand, remain saturated. Evaporation and consumption by soil organisms lower the moisture content in soils; at low moisture content, moisture exists as thin films and as wedges at the contact points of soil particles. Many of the chemical properties of water are attributed to its dipolar molecular structure. Water exhibits a high dielectric constant, which promotes the dissociation of many compounds in water. Thus, water is an excellent solvent for a number of chemical compounds. It is the most important transporting agent for nutrient elements and pollutants in soils. Many of the dissolved materials are in ionic forms. If a compound with a general formula A,B x is in contact with soil water, it dissociates into its ionic components An+ and B~-, where A is the metal cation with charge n and B is the anion with charge x. As in the case with protons (H+), which exist a s H3O+, a metal ion cannot exist by itself. In soil water, the metal reacts with water molecules, forming a hydration shell. The number of water molecules attracted to the metal ion depends on the coordination number of the cation. The hydrated cation carries the original number of positive charges, and is denoted A(H20)• n+. Water is held in the pore spaces by forces of attraction exerted by the soil matrix. These forces expressed in terms of matric, osmotic, and pressure potentials, are collectively known as soil water potential. The matric potential and the osmotic potential are the forces that bind the soil water to the soil solids and the soil solutes, respectively. The pressure potential, which results from pressure differences in soils, is responsible for the retention of water in soil pores (capillary forces) and on the surface of soil particles (adsorption). Several units have been used to express differences in energy levels of soil water. One is the height in centimetres of a unit water column whose weight just equals the potential under consideration. The pF unit is commonly used to characterize soil water potential. It is defined in

62

SOIL SYSTEM

terms of the logarithm of the height, in cm, of water, i.e., pF = lOgl0 cm H20. Practical values of soil water potential range from 1 cm of water (pF = 0) for saturated soil to 108 cm of water (pF = 7) for oven dry soil at 105-110~ A second is the standard atmosphere pressure at sea level, which is 760 mm Hg or 1020 cm of water. The unit termed bar approximates that of a standard atmosphere. At field capacity, the soil water potential is 0.3 bars, which corresponds to a pF value of 2.54. Energy may be expressed per unit of mass (joules/kg) or per unit of volume (newton/m2). Soil water potential has a negative value because of the matric and osmotic forces, which reduce the free energy level of the soil water. This means that soil water cannot move freely. The larger the negative value of soil water, the smaller the amount of water present in the soil. Generalized soil water potential curves for soils with different grain sizes are shown in Figure 4.1. Kaolinite clay holds much more water at a given potential level than does sandy clay or fine sand. Likewise, at a given moisture content, the water is held much more tenaciously in the kaolinite than in the other two soils. Soil texture clearly exerts a major influence on soil moisture retention. Soil structure also influences soil moisture-energy relationships. A well-granulated soil has more total pore space than one with poor granulation or one that has been compacted. The reduced pore space may result in a lower water-holding capacity. The compacted soil also may have a higher proportion of small- and medium-sized pores, which tend to hold water with greater tenacity than do larger pores.

Figure 4.1. Soil moisture variation with soil water potential for different soils.

The soil moisture-soil water potential relationship during drying differs from the relationship during wetting. This phenomena, known as hysteresis, is illustrated in Figure 4.1. Hysteresis is caused by a number of factors, including the nonuniformity of soil pores. As soils are wetted some of the smaller pores are bypassed, leaving entrapped air that prevents water penetration. Likewise, as a saturated soil dries, some of the macro-pores that may be surrounded only by micro-pores may

SOIL PHASES

63

not lose their water until the soil water potential is low enough to remove the water from the smaller pores. Soil water is generally classified into free water, capillary water, and hygroscopic water as shown in Figure 4.2. These classes are defined as: (1)free water: water held between the maximum retention capacity (pF 0) and the field capacity (pF 2.54). Saturated soil contains free water, (2) capillary water: water held between the field capacity (pF 2.54) and the hygroscopic water (pF 4.5). Hygroscopic water refers to the maximum amount of water adsorbed by soils from the atmosphere. Unsaturated soils contain capillary water, and (3) hygroscopic water: water held at pF 4.5.

Figure 4.2. Classification of soil moisture.

4.2.3

Solid Phase

The solid phase in soils consists of both inorganic and organic fractions. The inorganic fractions, derived from the weathering products of rocks, range in size from tiny colloids (< 2 ~tm) to large gravel and rocks (> 2 mm), and include many soil minerals, both primary and secondary. The inorganic fractions exert a tremendous effect on the physico-chemical properties of soils and the ability to retain chemicals. Separated according to size, the inorganic soil fraction can be divided into three major soil groups: sand, silt and clay. Sand grains are irregular in size and shape, and are not sticky and/or plastic when wet. Their presence in soil promotes a loose and friable condition which allows rapid water and air movement. They are chemically inert and do not carry electrical charges. Thus, they have low water-holding and cation exchange capacities. Silt particles are intermediate in size and possess characteristics between those of sand and clay. Some silt particles may be capped or coated with clay films as a result of weathering. Such particles may, therefore, exhibit some plasticity, stickiness, and adsorptive capacity for water and cations. Clay is the smallest particle in soil and has colloidal properties. It carries a negative charge and is chemically the most active inorganic soil

64

SOIL SYSTEM

constituent. The presence of clay contributes to: (1) high water-holding capacity, (2) large specific surface area, and (3) high cation exchange capacity. The inorganic soil fraction is composed of soil minerals, and is, therefore, also referred to as the mineral fraction of soils. Minerals are inorganic in nature, possessing definite physical characteristics and chemical compositions. Soil minerals can be grouped into primary and secondary minerals. Primary minerals are minerals that have been released by weathering from rocks in a condition that is chemically unchanged. These minerals constitute the sand fraction of soils. Secondary minerals are derived from the weathering of primary minerals. They are present in the clay fraction of soils. Organic components include plant and animal residues at various stages of decomposition, cells and tissues of soil organisms. Organic components, although normally present in much smaller quantities than inorganic components, may significantly alter a soil's properties.

4.3

MINERAL COMPOSITION

4.3.1

Primary and Secondary Minerals

The composition of soil minerals is variable and depends on the composition of the rocks from which they were derived. Rocks are mostly composed of the elements 02, Si, A1, Fe, Ca, Mg, Na, and K. These are, therefore, the elements usually found in soil minerals. Silicates and oxides are probably the most common soil minerals. The silicon (Si) in soil silicates is present in the form of silica tetrahedral, which constitute the basic units of the clay mineral. On the basis of the arrangement of the silica tetrahedral (SiO4) in the crystal structure, 15 common silicates, listed in Table 4.2, are formed. The first six silicates in Table 4.2 are primary minerals since they are typically inherited from parent material, as opposed to precipitated through the weathering process. The key structural entity in these minerals is the Si-O bond, which is a more covalent and, therefore, stronger bond than typical metal-oxygen bonds. The relative resistance of any one of the minerals to decomposition by weathering can be correlated positively with the Si/O molar ratio of its fundamental silicate structural unit. This is because a larger ratio means a lesser need to incorporate metal cations into the mineral structure for the purpose of neutralizing the oxygen anion charge. To the extent that metal cations are so excluded, the degree of co-valency in the overall bonding arrangement will be greater and the mineral will be more resistant to decomposition in the soil environment. For the first six silicates shown in Table 4.2, the Si/O molar ratios of their fundamental structural units are as follows: 0.5 (quartz and feldspar, SiO2) , 0.40 (mica, SizOs), 0.36 (amphibole, Si401~), 0.33 (pyroxene, SiO3), and 0.25 (olivine, SiO4). The decreasing order of the Si/O molar ratio is the same as the observed decreasing order of resistance to chemical weathering in the sand or silt fractions of soils. The minerals epidote, tourmaline, zircon, and rutile, shown in Table 4.2, are found to be highly resistant to weathering in soil environment. With reference to Table 4.2, listed minerals from kaolinite to gypsum are secondary minerals since they result from the weathering transformations of primary silicates. Often these secondary minerals are of clay size and exhibit poorly ordered atomic structure. Variability in their composition through the substitution of ions into their structure (isomorphous substitution) is also frequently noted in soils. The secondary silicates, smectite and vermiculite, bear a net charge on their surfaces,

MINERAL COMPOSITION

65

principally because of their variability in composition. Kaolinite and the secondary metal oxides below it in the list also bear a net surface charge due, however, to proton adsorption and desorption, not compositional variability. Metal oxides such as gibbsite and geothite tend to persist in the soil environment longer than the secondary silicates. This is because Si is more readily leached than A1, Fe, and Mn, unless significant amounts of soluble organic matter are present to render the metals more mobile.

Table 4.2: Common soil minerals (Bohn et al., 1979) Name

Chemical Formula

Importance Abundant in sand and silt Abundant in soil that is not leached extensively Source of K in most temperate-zone soils

Pyroxene Olivine

SiO 2 (Na,K)A102[SiO2]3 CaA1204[SiO212 KzA12OS[Si2OS]3A14(OH) 4 K2AlzOs[Si2Os]3(Mg, Fe)6(OH)4 (Ca, Na, K)z.3(Mg, Fe, A1)5(OH)2 [(Si, A1)40,112 (Ca, Na, Fe, Ti, A1)(Si, A1)O3 (Mg, Fe)2SiO4

Easily weathered to clay minerals and oxides Easily weathered Easily weathered

Epidote Tourmaline Zircon Rutile

Ca2(al, Fe)3(OH)Si3Ol2 NaMg3A16B3Si6027(OH, F)4 ZrSiO 4 TiO2

Highly resistant to chemical Highly resistant to chemical Highly resistant to chemical Highly resistant to chemical

kaolinite Smectite Vermiculite Chlorite

Si4A14Olo(OH)8 Mx(Si, A1)8(A1, Fe, Mg)4020(OH)4 Mx(Si, A1)8(A1, Fe, Mg)402o(OH)4 Mx(Si, A1)8(A1, Fe, Mg)402o(OH)4

Abundant Abundant Abundant Abundant

Allophane

Si3A14012.nH20

Imogolite

Si2A14Olo.5H20

Gibbsite Goethite Hematite Ferrihydrite Calcite Gypsum

AI(OH)3 FeO(OH) Fe203 FeloO15.9HzO CaCO 3 CaSO4.2H20

Abundant in soils derived from volcanic ash deposits Abundant in soils derived from volcanic ash deposits Abundant in leached soils Most abundant Fe oxide Abundant in warm regions Abundant in organic soils Most abundant carbonate Abundant in arid regions

Quartz Feldspar Mica Amphibole

Note: M represents interlayer cations.

weathering weathering weathering weathering

in clays due to weathering in clays due to weathering in clays due to weathering in clays due to weathering

66

SOIL SYSTEM

Organic matter is an important constituent of the solid fraction of soils. The structural complexity of soil organic compounds precludes the making of a simple list of component solids like that shown in Table 4.2. 4.3.2

Trace Elements in Soil Minerals

One of the most important aspects of the variability in composition of soil minerals is their content of trace elements. A trace element is any chemical element whose mass concentration in a solid phase is less than or equal to 100 mg/kg (Sposito, 1989). Soil minerals bearing trace elements serve as reservoirs for the elements, releasing them slowly into the soil solution as weathering of the minerals occurs. The bioavailability of an element depends on the rate at which it is transformed from a solid phase to a soluble chemical form. Soil physico-chemical properties such as pH, redox potential, and moisture content will affect the rate of this transformation and, thus, control element solubility. In this manner, the weathering rate of soil solids containing toxic elements (e.g., Cd) will determine, in part, the potential hazard to human, plants and animals. Trace elements in secondary soil minerals and soil organic matter are listed in Table 4.3. The chemical phenomenon underlying trace element occurrence is called coprecipitation, which is defined as the simultaneous precipitation of a chemical element with other elements by any mechanism and at any rate. The three broad types of coprecipitation are inclusion, adsorption, and solid-solution formation.

Table 4.3: Trace elements coprecipitated with secondary soil minerals and soil organic matter Solid Fe and A1 oxides Mn oxides Ca carbonates Illites Smectites Vermiculites Organic matter

Coprecipitated trace elements B, P, V, Mn, Ni, Cu, Zn, Mo, As, Se P, Fe, Co, Ni, Zn, Mo, As, Se, Pb P, V, Mn, Fe, Co, Cd B, V, Ni, Co, Cu, Zn, Mo, As, Se, Pb B, Ti, V, Cr, Mn, Fe, Co, Ni, Cu, Pb Ti, Mn, Fe A1, V, Cr, Mn, Fe, Ni, Cu, Zn, Cd, Pb

If a trace element forms a pure solid phase with atomic structure different from the host mineral, then two morphologically distinct solids will occur together. This kind of association is termed inclusion with respect to the trace element. For example, CuS often occurs as an inclusion, a small separate phase, in the primary silicates. If there is only limited structural compatibility between a trace element and the corresponding major element in a host mineral, coprecipitation can produce a homogeneous mixture of the two elements at the host mineral-soil solution interface. This mechanism is termed adsorption because the mixed solid phase is restricted to the interfacial region, which can change as the host mineral

SOIL M1NERAL TRANSFORMATIONS

67

continues to precipitate from the soil solution. Well known examples of adsorption are the incorporation of oxy-anions like borate, phosphate, or molybdate into secondary metal oxides, and of transition metals like Fe or Ni into soil organic matter. Finally, if structural compatibility is high and free diffusion of a trace element within the host mineral is possible, a major element in the host mineral can be replaced uniformly throughout by the trace element. This kind of homogeneous coprecipitation is called solid-solution formation. It is enhanced if the size and valence of the substituting element are comparable to those of the element replaced. Solid-solution formation occurs, for example, when secondary alumino-silicates precipitate and incorporate metals like Ni, Cu, and Zn to replace A1 in their structure or when calcium carbonate precipitates with Cd replacing Ca in the structure.

4.4

SOIL MINERAL TRANSFORMATIONS

The continual input and output of percolating water, biomass, and solar energy alters the composition of soils with the passage of time. The changes in clay fraction mineralogy observed during the course of soil development are shown in Table 4.4. These changes, known as JacksonSherman weathering stages, can be classified as early, intermediate, and advanced. Soils in stage 1 weathering may contain some gypsum and halite. Soils containing significant amounts of olivine are representative of stage 3, and biotite mica is representative of stage 4. This sequence is in agreement with their oxygen-silicon ratios and weathering resistance. Soils with minerals representative of stages 1 to 5 are considered to be in the early weathering stages. Such soils are often the least weathered, and are primarily found in regions where limited water restricts chemical weathering. Soils of the intermediate weathering stages (stages 6 to 9) include most soils of humid temperate regions. Quartz is often abundant in these soils and is representative of weathering stage 6. Hydrous mica, vermiculate, and montmorillonite are typically transformed in soils, and accumulate as fine-sized particles in the clay fraction. Soils of the advanced weathering stages include the intensely weathered soils of humid tropics. Soils dominated by minerals of weathering stages 10 to 13 may have lost all or most of the original minerals of the parent material. They may consist mainly of stable minerals that have been synthesized during weathering. These minerals are kaolinite, gibbsite, hematite, and anatase (TiO2).

4.5

CRYSTAL CHEMISTRY OF SILICATES

When atoms combine, a bond involving a redistribution of valence electrons is formed. The type of bond formed is a function of the electronic structure of the combining atoms. Ionic or electrostatic bonding occurs between ions of opposite charge, such as Na + and C1-. Such ions are formed by the complete loss or gain of electrons to form positive or negative ions having an electron structure like an inert gas. Ionic bonding forces are strong, and solid ionic compounds have high melting points. Ionic bonding forces are also undirected, that is, they are exerted uniformly in all directions. The valence of a given ion is shared by all neighboring ions of opposite charge. The number of such neighbors is determined by their size relative to the size of the central ion. Ionic bonds predominate in many inorganic crystals, including the silicate minerals.

68

SOIL SYSTEM

Table 4.4: Representative minerals associated with weathering stages (Jackson and Sherman, 1953) Weathering stage

Representative minerals

Typical soil group

EARLY WEATHERING STAGES Gypsum Soils dominated by these minerals, in Calcite, dolomite the fine silt and clay fractions, are Olivine-hornblende mainly weathered soils. They are of the Biotite desert regions, where limited water keeps Albite, microcline, orthoclase chemical weathering to a minimum.

INTERMEDIATE WEATHERING STAGES Quartz Soils dominated by these minerals are in Hydrous mica, muscovite the fine silt and clay fractions. They are Vermiculite mainly of temperate regions. Montmorillonite

10 11 12 13

ADVANCED WEATHERING STAGES Kaolinite Soils dominated by these minerals are in Gibbsite the clay fraction. They are intensely Hematite, geothite weathered soils of the warm and humid Analase, zircon equatorial regions. The soils are frequently characterized by acidity.

Covalent bond or shared electron pair bonding is common between identical atoms or atoms having similar electrical properties, such as H2O, F 2, and C H 4. Covalent bonding is the sharing of electron pairs between the combining atoms so that each atom attains the inert gas electronic structure. Covalent bonding is strong, but bonding is directional. Covalently bonded molecules have little tendency to ionize. Bonding within ionic radicals, such as SO42, is frequently covalent. Hydrogen bonding occurs between hydrogen and two atoms of high electro-negativity, such as F, O, and N. The H-bond is essentially a weak electrostatic bond but is nevertheless important in crystal structures of oxy-compounds, such as the layer silicates. Summed over many atoms, the individually weak H-bonds can strongly bond adjacent structures. van der Waals bonding is the weak electrostatic force between residual charges on molecules. Residual charges may result from natural dipoles of non-symmetrical molecules, polarization dipoles, or vibrational dipoles, van der Waals forces are generally obscured by stronger ionic and covalent bonding forces but may dominate the properties of some molecules. Although differences in the types of bonds described above are rather clear-cut, bonding in most crystals is not. For example, the Si-O bond in silicates is intermediate between the extremes of purely ionic and purely covalent bonding. The degree of ionic nature of the Si-O bond is

CRYSTAL CHEMISTRY OF SILICATES

69

sufficient, however, to apply the rules of ionic bonding to silicate structures. The internal bonding in silicates is predominantly ionic. As a result, forces are undirected and ionic size plays an important part in determining crystal structure. The crystal radii of common ions in the silicates are given in Table 4.5. The ionic radius of oxygen is much larger than that of most cations found in silicates. The oxygen ion constitutes nearly 50% of the mass, and over 90% of the volume, of most common silicate minerals. Hence silicate structures are largely determined by the manner in which the oxygen ions pack together. An ion in a crystal surrounds itself with ions of opposite charge. The number of ions that can be packed around a central ion depends on the ratio of radii of the two ions and is called the coordination number of the central ion. Ions are held together more or less rigidly in a crystal structure, determined by geometry and by electrical stability. More than one structure may meet the necessary requirements, but the most stable form will be the one in which the ions have the lowest potential energy. The requirement of electrical stability means that the sum of positive and negative charges must be equal. However, ions of opposite charge are not paired off to achieve neutrality. Instead, the positive charge of a cation may be considered to be divided equally among surrounding anions. The number of anions around each cation is determined by the coordination number, or radius ratio, of the cation and anions, rather than by the charge of the cation.

Table 4.5: Crystal ionic radii of selected cations, and their coordination number with oxygen (Bohn et al., 1979) Ion radius Ion Si 4+ A13+ F e 3+

Mg 2+ Fe 2§ Na § Ca > K+ NH4+ 02.

(A)

0.42 0.51 0.64 0.66 0.74 0.97 0.99 1.33 1.43 1.32

Coordination number with oxygen Observed Predicted 4 4,6 6 6 6 6,8 8 8,12 8,12 . . . . .

4 4 4 6 6 6 6 8 8

Predicted and observed coordination numbers of oxygen anion with common cations are given in Table 4.5 (Bohn et al., 1979). The Si 4§ cation occurs in fourfold or tetrahedral coordination. Aluminum is generally found in sixfold or octahedral coordination but may also occur in tetrahedral coordination. The tetrahedral and octahedral coordination units are basic to the atomic lattices of most layer silicate minerals. The tetrahedral structure, shown in Figure 4.3, consists of four 02. ligands coordinated around Si 4§ giving the ionic unit (SiO4) 4". The electrostatic bond strength (ion charge divided by number of bonds to the ion) for the tetrahedral unit is one. In fourfold

70

SOIL SYSTEM

coordination, the hole between the four O 2" ranges from 0.29 to 0.52 ,~. The radius of Si 4+, in fourfold coordination, is about 0.42 A, indicating some distortion from the ideal tetrahedral unit. The basic octahedral structure, shown in Figure 4.3, consists of six OH- groups coordinated around a central cation. The electrostatic bond strength is 1/2 or 1/3, depending on the charge of the central ion. Sixfold coordination yields an eight-faced structure, hence the name octahedral. The hole between O 2 ligands in this configuration has a theoretical radius of about 0.61 ,~. Ions commonly found in octahedral coordination in layer silicates include A13+ (radius 0.51 A in this coordination), Mg 2§ (0.66 ]k) and Fe z+ (0.74 A).

Figure 4.3. Silica tetrahedron and octahedrons of aluminum and magnesium.

4.6

STRUCTURAL COMPONENTS OF SOIL CLAYS

Many soil clays are structurally related. Thus, learning the basic components of soil clays facilitates an understanding of their nature and differences. Knowledge of clay structure is essential for understanding how clays affect the physical and chemical properties of soils, such as changes in hydraulic conductivity due to intrusion of pollutants, mobility of pollutants, soil adsorption capacity, buffering capacity, etc.

STRUCTURAL COMPONENTS OF SOIL CLAYS

4.6.1

71

Silica, Gibbsite, and Brucite Sheets

Many soil clays are alumino-silicates. As already noted, silicon-oxygen tetrahedra share oxygen atoms to form sheets. The apical oxygen has one excess negative charge, and the silica sheet has the formula Si2052-. The gibbsite sheet, shown in Figure 4.4, is composed of aluminum in six coordination with hydroxyl. An individual octahedron is AI(OH)63-. The sharing of hydroxyls by adjacent octahedra forms an octahedral or gibbsite sheet with the composition A12(OH)6. Two of the three potential spaces for aluminum in the sheet are filled, the mineral is referred to as dioctahedral. Although gibbsite sheet is an integral constituent of many alumino-silicate clays, it also exists by itself in soils.

Figure 4.4. Silica, gibbsite and brucite sheets.

9In a manner similar to aluminum and hydroxyl, magnesium and hydroxyl form an octahedral sheet called the brucite sheet, as shown in Figure 4.4. The magnesium is in six coordination with hydroxyl, and the sheet has the composition Mg3(OH)6. All three of the potential spaces for cations are filled with magnesium to produce a tri-octahedral mineral. Both gibbsite and brucite are octahedral and have electrically neutral structures. They differ in that gibbsite is dioctahedral and brucite is trioctahedral.

72

SOIL SYSTEM 1:1 LAYER SILICA TES

The structural unit of the kaolin group is formed by the superposition of a tetrahedral sheet upon an octahedral sheet. Such minerals are referred to as 1:1 layer silicates. The apical oxygens of the tetrahedral sheet are shared by the octahedral sheet, forming a common plane of oxygen ions within the structure, as shown in Figure 4.5. In the shared plane, two-thirds of the oxygen ions are shared between Si and A1. The remaining one-third of the oxygen ions have their charge satisfied by H § to form OH groups. The upper surface of kaolin is a layer of closely packed OH groups, but the bottom surface is composed ofhexagonally open-packed oxygens and OH groups recessed within hexagonal openings. Kaolinite [AI2Si2Os(OH)4] is the layer silicate mineral that represents the kaolin group. Silicon is apparently the only cation in the tetrahedral sheet of kaolinite, but A13§ or Mg 2+ may occupy the octahedral positions. If A13+is in octahedral coordination, the mineral is kaolinite or one of its poorly crystallized polymorphous forms (dickite or nacrite). With Mg 2+ in octahedral coordination, the mineral is antigorite [Mg3Si2Os(OH)4]. Halloydite is a form of kaolinite in which water is held between structural units in the basal plane, yielding a c-spacing of 10 ~, when fully hydrated. Most kaolin structural units, however, are held together in the basal plane by hydrogen bonding between oxygen ions of the tetrahedral sheet and hydroxyl ions of the octahedral sheet.

Figure 4.5. Schematic structure of kaolinite.

2:1 LA YER SILICA TES In 2:1 layer silicates, the unit layer is one octahedral sheet sandwiched between, and sharing oxygen atoms with, two tetrahedral sheets. The unit layers then stack parallel to each other in the cdimension. The atomic arrangement for illite mineral is shown in Figure 4.6.

STRUCTURAL COMPONENTS OF SOIL CLAYS

73

The various 2:1 minerals are differentiated by the kinds and amounts of isomorphous substitution in both the tetrahedral and octahedral sheets, which can lead to localized charge within the crystal. This excess charge must be balanced by other cations, either inside the crystal or outside the structural unit. The magnitude of charge per formula unit, when balanced by cations external to the unit layer, is called the layer charge. Typical layer charges of 2:1 minerals are shown in Table 4.6.

Figure 4.6. Schematic structure of illite.

The magnitude of a layer charge plays a dominant role in determining the strength and type of bonding in the basal plane. If the layer charge is zero, as in pyrophyllite, the basal planes of adjacent unit cells are bonded together by van der Waals forces. If the layer charge is negative, adjacent basal planes are bonded electrostatically by cations located between the unit layers. The greater the layer charge, the stronger the interlayer bond. Smectites of low layer charge have a weak bond, enabling polar molecules, such as water, to get between the basal planes and cause the minerals to expand. In minerals of high layer charge, such as the micas, the ionic bond is so strong that polar molecules cannot get between the basal planes, and the minerals are non-expanding. Vermiculites are intermediate in layer charge and also intermediate between mica and smectite in their expansion properties. Within a given mineral group, specific minerals are defined by the predominant ion in octahedral coordination, as shown in Table 4.6.

74

SOIL SYSTEM

Table 4.6: Classification of layer silicate minerals Layer charge per unit Mineral group Tetrahedral Octahedral Sheet sheet Pyrophyllite Talc Smectites Vermiculites Micas l

0 0.25 - 0.6 0 0.6 - 0.9 1

0 0 0.25 - 0.6 0 0

Octahedral cation A13+

Mg 2+

Pyrophyllite Beidellite Montmorillonite Vermiculite Muscovite

Talc Saponite Hectorite Vermiculite Biotite 1

Mg and Fe are in octahedral coordination; K is in the interlayer position.

In sepiolite and palygorskite, the 2:1 layers do not form continuous sheets, but form fibres six (sepiolite) or four (palygorskite) silicon tetrahedra wide. Simplified formulas are [Mg 4Si6 O~5(OH)2.6H20] and [(Mg, A1, Fe)4Si8 020. nH20], respectively. The symbolic structure of these minerals is shown in Figure 4.7. Sepiolite may occur as the pure Mg end member, but most sepiolites and all natural palygorskites contain some aluminum and usually some exchangeable cations.

Water k'.x'.'CZZZ].'.x'.'4 Water I f . f l Z Z Z b , ' . f . ~

Wa ter ~,xJ,,~,E~lYJ,~ Water ~ 7 , ] ~ ~ F 7 , 1

Water

Wa ter

Water rT~,eJ~,~,,~77~ Water

Figure 4.7. Symbolic structure ofpalygorskite and sepiolite. The 2:1 layers form chains rather than a continuous sheet.

2:1:1 LAYER SILICA TES The chlorite mineral group is closely related to the micas and has about the same layer charge. In chlorite, the interlayer potassium of mica is replaced by positively charged octahedral brucite [Mga(OH)6 ] sheet. The brucite sheet develops a positive charge when the Mg 2+ is partially replaced by A13+, yielding the basic unit [Mg2AI(OH)6] +1 that fits into the interlayer position of the

PROPERTIES OF LAYER SILICATES

75

2:1 layer silicates. This is referred to as a 2:1:1 type classification. Chlorites are non-expanding minerals with low cation exchange capacity.

4.7

PROPERTIES OF LAYER SILICATES

4.7.1

Kaolins

The kaolinite crystal consists of repeating layers, each layer consisting of a silica sheet and an alumina sheet sharing a layer of oxygen atoms between them, as shown in Figure 4.5. Each layer is three oxygen atoms thick. The layers are held together by hydrogen bonding between hydroxyls from the alumina sheet on one face and oxygens from the silica sheet on the opposite face. These forces are relatively strong, preventing hydration between layers and allowing many layers to build up. A typical kaolinite crystal may be between 70 to 100 layers thick. Kaolinite occurs commonly in soils, often as hexagonal crystals with an effective diameter of 0.2 to 2 ~tm. Hydrogen bonding between adjacent unit layers prevents expansion of the mineral beyond its basal spacing of 7.2 A. Surface area is limited to external surfaces and, hence, is relatively small, ranging from 10 to 20 m2/g. Kaolinite is a coarse clay with low colloidal activity, that is, low plasticity and cohesion, and low swelling and shrinkage.

Table 4.7: Summary of selected properties of solid phase components Cation Mineral Layer Exchange Surface Component type charge Capacity area (meq/100~) (m2/~) Kaolinite Montmorillonite Vermiculite Mica Chlorite Organic matter

1:1 2:1 2:1 2:1 2:1:1 ..........

0 0.25-0.6 0.6-0.9 1.0 -~1.0

1-10 80-120 120-150 20-40 20-40 100-300

10-20 600-800 600-800 70-150 70-150 800-900

c-spacing

A 7.2 variable 10-15 10 14 ....

The ideal unit formula for kaolinite [A12Si2Os(OH)4] has an Si/A1 ratio of one which suggests little or no isomorphous substitution. Most of the cation exchange capacity (1 to 10 meq/100g) of kaolinite can be attributed to the dissociation of OH groups on clay edges. The cation exchange capacity of kaolinite is highly pH-dependent, suggesting that isomorphic substitution is not the predominant source of charge. A summary of some selected properties of kaolinite is shown in Table 4.7.

76

4.7.2

SOIL SYSTEM

Hydrous Mica (Illite)

Mica minerals have repeating layers of an alumina sheet between two silica sheets, with shared oxygen to give a unit of four oxygen atoms thick. The layers are bonded together by potassium ions which are just the right size to fit into the hexagonal holes of the silica sheet, as shown in Figure 4.6. The bonding, via potassium ions, between adjacent unit layers prevents expansion of the mineral beyond its basal spacing of 10 A. The potassium ions exist in twelve coordination, bonding six oxygens from one silica sheet to the adjacent six oxygens of the silica sheet of the next layer. Negative charge, to balance the potassium cations, arises from the substitution of aluminum for silicon in the silica sheet (isomorphous substitution). A typical unit formula for mica is K[AI2(Si3A1)OI0(OH)2]. Despite the relatively large layer charge (-- -1) of the mica, its cation exchange capacity is only 20 to 40 meq/100g. Its total surface area is about 70 to 120 m2/g and is restricted to external surfaces, as indicated in Table 4.7. Soils containing illite have properties intermediate between kaolinite (low activity) and montmorillonite (high activity). Illite occurs widely in temperate and in arid regions.

Figure 4.8. Schematic representation of montmorillonite structure.

4.7.3

Montmorillonite

A typical unit formula is Nax[(A12_xMgx)Si40~0(OH)2] in which Na § is the chargecompensating exchangeable cation. Montmorillonite minerals have the same layers as micas, discussed in the previous section. However, soil montmorillonites exhibit imperfect isomorphic

PROPERTIES OF LAYER SILICATES

77

substitution, with some A13+ substituting for Si4+ in the tetrahedral sheet and with Fe 2+(as well as Mg 2+) substituting for A13+ in the octahedral sheet. There are no potassium ions to bond the layers together, and water enters easily between layers, as illustrated in Figure 4.8. The distance of separation of the layers on hydration can be controlled if certain organic liquids rather than water are used. Montmorillonite saturated with glycerol will show layer spacings of 17.7 ~,, of which 10 A is the thickness of the layer and 7.7 A of the glycerol. Typical cation exchange capacities for montmorillonite range from 80 to 120 meq/100g, as indicated in Table 4.7. The cation exchange capacity is only slightly pH-dependent. The lower layer charge allows the mineral to expand freely, exposing both internal and external surfaces. Such expansion yields a total surface area of 600 to 800 m2/g, with as much as 80% of the total surface area due to internal surfaces. Montmorillonite has high colloidal activity, that is, high plasticity and cohesion, and high swelling and shrinkage. Montmorillonite normally occurs as a fine clay with irregular crystals having an effective diameter of 0.01 to 1 gm. The combination of high specific surface area, cation exchange capacity and swelling potential of montmorillonite makes it attractive for use as a waste barrier material. The interlayer spacing, which can include several water layers, will, however, not respond in a similar fashion in the presence of certain organic pollutants. In essence, when water saturated montmorillonite is exposed to certain organic pollutants, penetration of the organic pollutants, known as intercalation phenomenon, into the saturated montmorillonite occurs easily. 4.7.4

Vermiculites

Vermiculites occur extensively in soils formed as a product of weathering or hydrothermal alteration of micas. The layer structure of vermiculite resembles that of the mica from which the mineral is derived. Due to weathering, the interlayer K + in the micas is replaced by Mg 2+, and the c-spacing expands, in most cases, to 14-15 ,&. An idealized unit formula is [Mg(H20)6]n[(Mg , Fe)3(Si4_n, A1,)O10(OH)2], with the hydrated magnesium cation Mg(H20)62+ serving as the exchangeable cation. The layer charge in vermiculite gives rise to a cation exchange capacity of 120 to 150 meq/100g, which is considerably higher than the exchange capacity of montmorillonite. As with montmorillonite, the cation exchange capacity is only slightly pH-dependent. Vermiculite swells less than montmorillonite because of its higher layer charge. The total surface area of vermiculite ranges from 600 to 800 m2/g. The mineral, with a basal spacing of 10 A, is non-swelling when saturated with K + or NH4+ ions. A summary of some selected properties of vermiculite is shown in Table 4.7.

4.7.5

Chlorites

Chlorites occur extensively in soils and are examples of 2" 1"1 layer silicates. The positively charged mica-like sheet restricts swelling, decreases the effective surface area, and reduces the effective cation exchange capacity of the mineral. An idealized unit formula is [A1 Mg2(OH)6]x [Mg3(Si4.xAlx) O10(OH)2]. Substitution in such classical chlorites is in the tetrahedral layer, with the brucite sheet serving as the interlayer cation, as illustrated in Figure 4.9. The repeating layer has a thickness of 14 A. The layer charge of the 2"1 portion of the mineral is variable but is similar to that of mica. Cation exchange capacity ranges from 10 to 40 meq/100g, and total surface area from 20

78

SOIL SYSTEM

to 150 m2/g. A summary of some selected properties of chlorites is shown in Table 4.7.

si

/

Si

"ZJ _Y

si

/

Si

Figure 4.9. Schematic representation of typical chlorite structure.

4.7.6

Sepiolite and Palygorskite

In sepiolite and palygorskite, the 2:1 layers do not form continuous sheets, but form fibers six (sepiolite) or four (palygorskite) silicon tetrahedra wide, as shown in Figure 4.7. Simplified formulae are [Mg 4Si60!5(OH)2.6H20] and [(Mg, A1, Fe)4Si8020. nH20], respectively. Sepiolite may occur as the pure Mg end member, but most sepiolites and all natural palygorskites contain some aluminum and usually some exchangeable cations. Cation exchange capacity ranges from 20 to 30 meq/100g, and total surface area from 170 to 370 m2/g.

4.7.7

Mixed-layer Clays

The structures of2:1 clays and chlorites are closely related. It is not surprising, therefore, that same minerals contain more than one type of interlayer behaviour. Some layers in a crystal may be of the smectite type and some of the illite type, giving a mixed-layer illite-smectite. Also, regions of gibbsite or brucite may occur between the layers of a smectite or vermiculite, giving what is known as a mixed-layer chlorite-smectite or a hydroxy-interlayer smectite. The different layers may be distributed randomly or may exhibit several types of ordering, making precise identification of mixed-layer structure difficult.

4.7.8

Soil Clays

As mentioned previously, soil clays often differ appreciably in properties from those of the pure minerals described above. Soil clays are usually less well ordered and smaller in size than the pure minerals and often overlap neighboring particles or sheets. Inter-stratifications of various layer silicates are common, and the mineralogy of soil clays is rarely simple or uniform. Coating of iron and aluminum oxides and organic matter on most layer silicates further complicate the mineralogy

SOIL ORGANIC MATTER

79

of soil clays. Such coatings can drastically alter mineral properties by decreasing cation exchange capacity and surface area values and by restricting the swelling and collapsing of expansible minerals. Oxide coatings, however, magnify anion exchange and other properties associated with positively charges surfaces.

4.8

SOIL ORGANIC MATTER

Soil organic matter is an accumulation of partially decayed and partially re-synthesized plant and animal residues. Such material is in an active state of decay, being subjected to continued attack by soil microorganisms. Consequently, much of it is rather transitory and must be constantly renewed by addition of plant residues. The organic matter content of surface mineral soils is usually only about 0.5 to 5% by weight. Soil organic matter can exert a profound effect on the physical and chemical properties of the soil. Physically, it improves aggregation of soil particles, resulting in the development of a stable soil structure. Chemically, it increases the cation exchange capacity, and the water holding capacity of soils. Biologically, soil organic matter is the main source of food and energy for soil organisms. The accumulation of organic matter in soil is strongly influenced by temperature and the availability of oxygen. Since the rate of biodegradation decreases with decreasing temperature, organic matter does not degrade rapidly in colder climates and tends to build up in soil. In water and in waterlogged soils, decaying vegetation does not have easy access to oxygen, and organic matter accumulates. Non-humus organic matter includes those materials that are undecomposed (original tissue) or only partially decomposed. Non-humus substances include carbohydrates and related compounds, proteins and their derivatives, fats, lignins, tannis, and various decomposition products. Non-humus organic matter may also include roots and tops of plants. The degradation products of non-humus materials undergo enzymatic and chemical reactions to form new colloidal polymers called humus. Humus is a generic term for the water-insoluble material that makes up the bulk of soil organic matter (Stevenson, 1994). Humus is composed of a base-soluble fraction (humic and fulvic acids) and an insoluble fraction (humin), and is the residue from the biodegradation (by bacteria and fungi) of plant material. The bulk of plant biomass consists of relatively degradable cellulose and degradation-resistant lignin, a complex polymeric substance that is second only to carbohydrates in natural abundance (Sarkanen and Ludwig, 1971). Humic materials in soil strongly adsorb many solutes in soil water and have a particular affinity for polyvalent catioo~ and interact with the clay minerals. Both the humus and non-humus fractions of soil organic matter are important to the soil environment. Non-humus material provides short-range effects, such as sources of food and energy for microorganisms. Humus provides long-term effects, such as maintaining good soil structure and increasing soil cation exchange, pH-buffering, and water holding capacity. Thus, humus reduces bioavailability.

80 4.9

SOIL SYSTEM CHARGE DEVELOPMENT IN SOILS

The two properties that most account for the reactivity of soils are surface area and surface charge. Surface area is a direct result of particle size and shape. Most of the total surface area of a mineral soil is due to clay size particles and soil organic matter. Charge development in soils is associated with these two fractions, although the sand and silt size fractions may contribute some cation exchange capacity if coarse grained vermiculite is present. A charge develops in soils through isomorphic substitution and ionization of functional groups on the surface of solids that make up the soil matrix. These two mechanisms give rise to the constant surface charge minerals and the constant surface potential (pH dependent charge) clay minerals. The separation is not a rigid one because a single soil mineral can exhibit both types. 4.9.1

Constant Surface Charge Minerals

A perfectly formed crystal lattice would possess no excess charge at the surface because all atoms in the crystal would be electrically balanced. Imperfections in the lattice structure, however, cause an excess of positive or negative charge, which is then compensated for by the accumulation of oppositely charged ions (counter ions) at the crystal surfaces. Such an imperfection, for instance, might be the substitution of the trivalent aluminum atom in a silicate sheet, which would lead to an excess of negative charge at the particle surface. The substitution of trivalent aluminum for divalent magnesium would lead to an excess of positive charge at the surface. This type of substitution is called isomorphous substitution. As this defect occurs in the interior of the crystal lattice, the resulting charge imbalance is permanent and cannot be influenced by external factors such as the pH of the ambient solution. Hence, we have a constant surface charge mineral. 4.9.2

Constant Surface Potential Minerals

In this general type, surface charge is created by the adsorption of ions onto the surface, the net charge being determined by that ion which is adsorbed in excess. The charging process requires the presence of these ions, called potential determining ions, in the ambient solution in quantities sufficient for adsorption. The primary source ofpH-dependent charge is considered to be the gain or loss ofH + from functional groups on the surfaces of soil solids. The functional groups include hydroxyl [-OH], carboxyl [-COOH], phenolic [-C6H4OH], and amine [-NH2]. The charge that develops from functional groups depends largely on the pH of the ambient solution, which regulates the degree of protonation or deprotonation of the functional group. The soil solids that contain functional groups capable of developing pH-dependent charge include layer silicates, oxides and hydrous oxides, and soil organic matter.

Protonation of Exposed OH Groups Exposed OH groups are present on the surface of A1 octahedral sheets. They are prevalent in 1:1 types of clays, oxides, and amorphous soils. These OH groups are in contact with the soil solution and tend to protonate due to the addition of H+ ions. This process contributes to the oversaturation of the OH groups with protons, thus rendering the clay surface positively charged, as

CHARGE DEVELOPMENT IN SOILS

81

illustrated by the following reaction: - A I - O H + H* -~ - A I - O H H + neutral positively charged octahedral octahedron

[4.2]

Protonation of exposed OH groups occurs only at low pH, because acid conditions are required for the supply of the extra proton. The positively charged octahedron will contribute to the increased mobility of cations through the soils, hence increasing the potential of polluting the groundwater.

(-1) OH

/

/

Si (+1/2)

AI

(+1) ~

~ +

H

/

H

Acid

Si

(+1/2) OH

~ +2OH" ~

AI

/ (+ 1/2) OH

0

/

Si OH

/

OH

(-1/2) 0 § 2H20

/ AI

(-1/2) OH (-1/2) Neutral pH

OH Basic

Figure 4.10. Representation of pH-dependent charge at kaolinite edges.

Deprotonation o f Exposed O H Groups

Since OH groups are in contact with soil water, they tend to dissociate (deprotonate), and release their protons, as illustrated in the following reaction: - A I - O H -~ - A I - O - + H + neutral negatively charged octahedral octahedron

[4.3]

The dissociation of H + leaves one non-neutralized negative charge in the octahedron. Such a dissociation reaction occurs at high pH, and decreases at low pH. The magnitude of the negative charge also increases and decreases accordingly with the change of pH. Therefore, this type of

82

SOIL SYSTEM

negative charge is called pH-dependent or variable charge. Figure 4.10 illustrates the pH-dependent charge at kaolinite edges. The acidity of OH groups can be characterized by using the dissociation constant, pKa, with a value of 5.0 assigned to the AI(OH2)+i group, 7.0 to the (A1-OH-Si) +~ group, and 9.5 to the SiOH group (Sposito, 1989). The high pKa value for SiOH groups indicate that their deprotonation occurs only at high pH. Thus, variations in pH-dependent charge of layer silicates are more likely associated with reversible protonation and deprotonation of exposed A1OH groups. Hydroxyl ions exposed on planar surfaces of minerals are also characterized by high pKa values and contribute pH-dependent charge only at high pH. pH-dependent charges are more important for kaolinite than for smectites, illite and vermiculites. As a rule of thumb, only 5 to 10% of the negative charge on 2:1 layer silicates is pHdependent, whereas 50% or more of the charge developed on 1:1 minerals can be pH-dependent.

Zero Point of Charge The zero point of charge (ZPC) is the pH at which a mineral has no charge, or has equal amounts of negative and positive charges. It is similar in meaning to the isoelectric point. As previously discussed, at high pH values the mineral carries a negative charge, which decreases with a decrease in pH. When the pH is continuously decreased, a point will be reached at which the negative charge equals zero. The pH at which this occurs is the ZPC. Typical values of ZPC, of selected minerals, are shown in Figure 4.11. The ZPC is a specific characteristic of the clay mineral, and its value differs from one mineral to another. When the net charge is zero, at the ZPC, clay particles in soil water will not repel each other but will tend to aggregate and form larger particles. This in turn will contribute to an increase in soil hydraulic conductivity and transport of pollutants through soils. In contrast, negatively charged clay particles repel each other, resulting in dispersion and a decrease in soil hydraulic conductivity.

4.10

SURFACE FUNCTIONAL GROUPS The common surface functional groups on inorganic solids are discussed below (Sposito,

1989).

Lewis Acid Site The combination of metal cation and water molecule at an interface is a Lewis acid site, with the metal cation identified as the Lewis acid. For example, at the periphery of gibbsite mineral, water molecules are bound to A13+ions, which result in a positive charge. Lewis acid sites can exist also on the surface of geothite if peripheral Fe3+ions are botmd to water molecules there. Thus, any metal hydrous oxide, as well as the edge surfaces of clay minerals like kaolinite, can expose Lewis acid sites to the soil solution. These surface functional groups are very reactive, since the positively charged water molecule is unstable and is exchanged readily for an organic or inorganic anion in the soil solution, which then can form a more stable bond with the metal cation.

SURFACE FUNCTIONAL GROUPS

83

Hydroxyl Group The inorganic surface functional group of greatest abundance and reactivity in soil clays is a hydroxyl group that is exposed on the outer periphery of a mineral. This kind of OH group is found on metal oxides, oxy-hydroxides, and hydroxides on clay mineral and on amorphous silicate minerals like allophane. In the case of soil organic matter, the surface functional groups are organic molecular units. But in general they can bound to either organic or inorganic solids, and they can have any molecular structural arrangement. The main functional groups are hydroxyl [-OH], carboxyl [-COOH], phenolic [-C6H4OH], and amine [-NH2].

Hematite

Kaolinite

Amorphous Iron Gibbsite

to to

2.1

1

2

.

I

1

3

4

8.5~

J

5

6

I

7

8

pH Figure 4.11. Zero point of charge (ZPC) values of selected minerals.

4.11

S U M M A R Y AND CONCLUDING R E M A R K S

Soil is a multi-component system consisting of solid, liquid, and gaseous phases, and living organisms. The solid phase of soils consists of both inorganic and organic components. Inorganic components exert a tremendous effect on the physical and chemical properties, such as cation exchange capacity and surface area, and on the overall suitability of soil as a barrier for waste containment. The organic components, although normally present in much smaller quantities than inorganic components, may significantly alter soil properties. The variability of these separate soil components and pore fluid chemistry will impact on the nature of solid-pore fluid interaction mechanisms, adsorption capacity, and fluid transport properties such as hydraulic conductivity, diffusion and dispersion. These mechanisms and properties are important in evaluating the fate of chemical substances in the terrestrial ecosystem and determining the proper clay mixture for designing waste containment barrier systems. From a soil cleanup viewpoint, evaluation of the effectiveness of a decontamination procedure can be achieved from a closer consideration of how the pollutants are retained in the organic and inorganic solid phases.