Exchangeable Ions

In Section 3.4, the cation exchange capacity, or CEC, of particulate soil humus is defined as the maximum number of moles of proton charge per kilogram that can be desorbed by a metal cation under prescribed conditions. Thus, CEC for particulate humus is equal to the maximum absolute value of the negative net proton charge. Operationally, this maximum value is measured typically as the surface excess of Ba2+ adsorbed by humus at pH 8.2 (Eq. 3.5). Extending this concept to soils, one can define the CEC as the maximum number of moles of readily exchangeablemetal cation charge per unit mass of dry soil that can be extracted under prescribed conditions. In this more general context, CEC refers to metal cations that adsorb on soil particles in either outer sphere surface complexes or the diffuse ion swarm (Fig. 7.2). In alkaline soils, the common readily exchangeable cations are Ca2+, Mg2+, Na+, and K+, whereas in acidic soils, this group expands to include Al3+, and its complexes AlOH2+, Al(OH)2+, and AlSO+4. Following the operational paradigm for soil humus, one concludes that the measurement of soil CEC involves not only the desorption of protons, but also the replacement of the population of readily exchangeable adsorbed metal cations at a selected pH value (usually pH 7–8) by a chosen cation. Laboratory procedures for measuring CEC are described in Methods of Soil Analysis, listed in For Further Reading at the end of this chapter. In alkaline soils, the replacing cation chosen is often Na+ or Ca2+, whereas in acidic soils and for soil humus, the replacing cation of choice is Ba2+. These cations, in turn, are typically displaced from soil particle surfaces by Mg2+ to measure the surface excess. A conceptual definition of CEC can be developed in terms of the surface charge balance concepts introduced in Chapter 7. Consider first a soil in which a net positive surface excess of anions does not occur, such as the Mollisol example discussed in Section 8.1. In this case, the only adsorbed ions are Ca2+ and Cl-. The CEC of this soil may be defined by a special case of the charge-balance condition in Eq. 7.3a: ∆qex (max) ≡ CEC

Soil Particle Surface Charge

Structural charge arises on the surfaces of soil mineral particles in which either cation vacancies or isomorphic substitutions of cations by cations of lower valence occur. The principal minerals bearing structural charge are therefore the micas (Section 2.2), the 2:1 clay minerals (Section 2.3), or the Mn(IV) oxide, birnessite (Section 2.4). These three classes of mineral are all layer type and the cleavage surface on which their structural charge is manifest is a plane of O ions. The plane of O ions on the cleavage surface of a layer-type aluminosilicate is called a siloxane surface.This plane is characterized by hexagonal symmetry in the configuration of its constituent O ions, as shown at the top of Fig. 2.3 and, more explicitly, on the right side of Fig. 2.4, where a portion of the siloxane surface of the micas is depicted. Reactive molecular units on the surfaces of soil particles are termed surface functional groups. The functional group associated with the siloxane surface is the roughly hexagonal (strictly speaking, ditrigonalbecause the hexagonal symmetry is distorted when the tetrahedral sheet is fused to an octahedral sheet to form a layer) cavity formed by six corner-sharing silica tetrahedra. This cavity has a diameter of about 0.26 nm. The reactivity of the siloxane cavity depends on the nature of the electronic charge distribution in the layer structure. If there are no nearby isomorphic cations substitutions to create a negative charge, the O ions bordering the siloxane cavity function as an electron cloud donor that can bind molecules weakly through the van der Waals interaction. These interactions are akin to those underlying the hydrophobic interaction, discussed in Section 3.5, because the O in the siloxane surface can form only very weak hydrogen bonds with water molecules. Therefore, uncharged patches on siloxane surfaces may be considered hydrophobic regions to a certain degree, with, accordingly, an attraction for hydrophobic organic molecules. However, if isomorphic substitution of Al3+ by either Fe2+ or Mg2+ occurs in the octahedral sheet, the resulting structural charge is manifest on the siloxane cavities, as discussed in Section 2.3.

Soil Acidity

A soil is acidic if the pH value of the soil solution is less than 7.0. This condition is met in many soils where rainfall exceeds evapotranspiration, including Alfisols, Histosols, Inceptisols, Oxisols, Spodosols, and Ultisols—almost half of the ice-free land area worldwide. Soils of the humid tropics offer examples of acidic soils (Ultisols and Oxisols), as do soils of forested regions in the temperate zones of Earth (Alfisols, Histosols, Inceptisols, and Spodosols). Soils in peat-producing wetlands and those influenced strongly by oxidation reactions, such as rice-producing uplands, can be mentioned as examples in which the biota play a direct role in acidification. The phenomena that produce a given proton concentration in the soil solution to render it acidic are complex and interrelated. Those pertaining to sources and sinks for protons are shown in Fig. 11.1, which is a special case of Fig. 1.4 with “free cation or anion” in the center of the latter figure now interpreted as H+. In addition to the biogeochemical determinants of soil acidity, the field-scale transport processes wetfall (rain, snow, throughfall), dryfall (deposited solid particles), and interflow (lateral movement of soil water beneath the land surface down hill slopes) carry protons into a soil solution from external sources. Their existence and that of proton-exporting processes, such as volatilization and erosion, underscore the fact that the soil solution is an open natural water system subject to anthropogenic inputs that may dominate the development of soil acidity. Industrial effluents, such as sulfur and nitrogen oxide gases or mining waste waters, that produce acidic deposition or infiltration, and nitrog-enous fertilizers, the transformation and transport of which produce acidic soil conditions, are examples of anthropogenic inputs. Despite all this complexity, proton cycling in acidic soils at field scales has been quantified well enough to allow some general conclusions to be drawn. Acidic deposition, production of CO2(g) and humus, plus proton biocycling, all serve to increase soil solution acidity, whereas proton adsorption and mineral weathering serve to decrease it.

Soil Adsorption Phenomena

Adsorption experiments involving soil particles typically are performed in a sequence of three steps: (1) reactio of an adsorptive (ion or molecule) with a soil contacting an aqueous solution of known composition under controlled temperature and applied pressure for a prescribed period of time, (2) separationof the wet soil slurry from the supernatant aqueous solution, and (3) quantitationof the ion or molecule of interest, both in the aqueous solution and in the separated soil slurry along with its entrained soil solution. The reaction step can be performed in either a closed system (batch reactor) or an open system (flow-through reactor), and it can proceed over a time period that is either relatively short (to investigate adsorption kinetics) or very long (to investigate adsorption equilibration). The separation step is similarly open to choice, with centrifugation, filtration, or gravitational settling being conventional methods to achieve separation. The quantitation step, in principle, should be designed not only to determine the moles of adsorbate and unreacted adsorptive, but also to verify whether unwanted side reactions, such as precipitation of the adsorptive or dissolution of the adsorbent, have influenced the experiment. After reaction between an adsorptive i and a soil adsorbent, the moles of i adsorbed per kilogram of dry soil is calculated with the standard equation ni ≡ niT − Mwmi where niT is the total moles of species i per kilogram dry soil in a slurry (batch process) or a soil column (flow-through process), Mw is the gravimetric water content of the slurry or soil column (measured in kilograms water per kilogram dry soil), and mi is the molality (moles per kilogram water) of species i in the supernatant solution (batch process) or effluent solution (flow-through process). Equation 8.1 defines the surface exces, ni, of an ion or molecule adsorptive that has become an adsorbate. Formally, ni is the excess number of moles of i per kilogram soil relative to its molality in the supernatant solution. As mentioned in Section 7.2, this surface excess may be a positive, zero, or negative quantity.

Oxidation– Reduction Reactions

Soils become flooded occasionally by intense rainfall or by runoff, and a significant portion of soils globally underlies highly productive wetlands ecosystems that are inundated intermittently or permanently. Peat-producing wetlands (bogs and fens) account for about half the inundated soils, with swamps and rice fields each accounting for about one-sixth. Wetlands soils hold about one-third of the total nonfossil fuel organic C stored below the land surface, which is about the same amount of C as found in the atmosphere or in the terrestrial biosphere. This C storage is all the more impressive given that wetlands cover less than 6% of the global land area. On the other hand, wetlands ecosystems are also significant locales for greenhouse gas production. They constitute the largest single source of CH4 entering the atmosphere, emitting about one-third the global total, with half this amount plus more than half the global N2O emissions coming from just three rice-producing countries. A soil inundated by water cannot exchange O2 readily with the atmosphere. Therefore, consumption of O2 and the accumulation of CO2 ensue as a result of soil respiration. If sufficient humus metabolized readily by the soil microbiome (“labile humus”) is available, O2 disappearance after inundation is followed by a characteristic sequence of additional chemical transformations. This sequence is illustrated in Fig. 6.1 for two agricultural soils: a German Inceptisol under cereal cultivation and a Philippines Vertisol under paddy rice cultivation. In the German soil, which was always well aerated prior to its sudden inundation, NO3- is observed to disappear from the soil solution, after which soluble Mn(II) and Fe(II) begin to appear, whereas soluble SO42- is depleted (left side of Fig. 6.1). The appearance of the two soluble metals results from the dissolution of oxyhydroxide minerals (Section 2.4). Despite no previous history of inundation, CH4 accumulation in the soil occurs and increases rapidly after SO42- becomes undetectable and soluble Mn(II) and Fe(II) levels have become stabilized. During the incubation time of about 40 days, the pH value in the soil solution increased from 6.3 to 7.5, whereas acetic acid (Section 3.1) as well as H2 gas were produced.

Mineral Weathering

Mineral weathering begins with mineral dissolution, typically as induced by protons or by ligands that form strong complexes with metals (Section 1.4). Proton-induced dissolution begins with H+ adsorption, exemplified in Eq. 3.2 for a metal oxyhydroxide mineral. In the absence of ligands that could replace the positively charged water molecule resulting from this rapid reaction, proton adsorption is followed by slow detachment of the metal, which then equilibrates as a soluble species in the soil solution, as illustrated in Fig. 5.1 (pathway 1) for gibbsite [Al(OH)3; see Fig. 2.7] at a pH value low enough that the detached Al3+ does not hydrolyze. Ligand-induced dissolution is also illustrated in Fig. 5.1 (pathway 2). The ligand is a fluoride anion, which forms a strong complex with Al3+ (see problem 3 in Chapter 4). Adsorption in this case occurs by ligand exchange, which is illustrated for carboxylate in Eq. 3.3. A similar reaction occurs for F-:...Slow detachment of the AlF2+ complex then follows. Whenever a mineral dissolution reaction induced by either of these two-step mechanisms is far from equilibrium, it is not influenced by the very low concentration of the constituent released from the dissolving mineral and its rate can be described by zero-order kinetics (Table 4.2). Accordingly, if [A] is the concentration of a constituent released, then the rate law can be expressed as...where kd is a rate coefficient independent of [A] , but a function of temperature, pressure, pH, the chemical properties of the mineral, and, if appropriate, the concentration of the ligand inducing dissolution via the second mechanism in Fig. 5.1. The mineral dissolution rate on the left side of Eq. 5.2 can be mass-normalized to express it in moles per mole of mineral per second by dividing the molar concentration [A] with the solids concentration of the mineral expressed in units of moles per liter. This mass-normalized rate does not depend on the amount of mineral dissolving. For proton-induced dissolution, the rate is then a function of temperature, pressure, pH, and the chemical nature of the mineral.

Soil Salinity

A soil is salineif the electrical conductivity of its soil solution as obtained by extraction from a water-saturated soil paste (ECe) exceeds 4 dS m-1. (The measurement of electrical conductivity for a soil saturation extract is discussed in Methods of Soil Analysis,listed under For Further Reading at the end of this chapter.) According to this definition, about a quarter of the agricultural soils worldwide are saline, but values of ECe > 1 dS m-1 are encountered typically in arid-zone soils, which cover almost one-third of the global ice-free land area. Ions released into the soil solution by mineral weathering, or introduced there by the intrusion of saline surface water or groundwater, tend to accumulate in the secondary minerals formed as the soils dry. These secondary minerals typically include clay minerals (Section 2.3), carbonates and sulfates (Section 2.5), and chlorides. Because Na, K, Ca, and Mg are brought into the soil solution relatively easily—either as displaced exchangeable cations or as cations dissolved from carbonates, sulfates, and chlorides—it is this set of four metals that contributes most to soil salinity. The corresponding set of anions that contributes to salinity is CO3, SO4, and Cl. Thus, arid-zone soil solutions are essentially electrolyte solutions containing chloride, sulfate, and carbonate salts of four metal cations. According to Eq. 4.21, an electrical conductivity of 4 dS m-1 corresponds to an ionic strength of 58 mM (log I = -1.841 + 1.009 log4 = 0.0584). This level of salinity is less than 10% of that of seawater (EC = 46.21 dS m-1), but high enough that only crops that are relatively salt tolerant can withstand it. Moderately salt-sensitive crops are affected when the electrical conductivity of a soil saturation extract approaches 2 dS m-1, corresponding to an ionic strength of 29 mM, and salt-sensitive crops are affected at 1 dS m-1 (I = 14 mM). Thus, with respect to crop salinity tolerance, a soil can be judged saline at any saturation extract ionic strength greater than 15 mM if crops are stressed.

Soil Colloids

Soil colloidsare solid soil particles with diameters ranging from 0.01 to 10 μm, which means they range from clay to fine silt in size. The chemical composition of these particles may be that of a single mineral or humus, but usually they are heterogeneous mixtures of inorganic and organic materials. Regardless of their composition, the characteristic properties of soil colloids are that they are small in size and relatively insoluble in water. Soil colloids exhibit shapes and sizes that reflect both chemical composition and the effects of weathering processes. Kaolinite particles, for example, are roughly hexagonal plates comprising perhaps 50 unit layers, with each unit layer being a wafer having the thickness of about 0.7 nm, which are stacked irregularly and held together through hydrogen bonding. In soils, weathering produces rounding of the corners of the kaolinite hexagons and coats them with iron oxyhydroxide and humus polymers (Fig. 10.1). Fracturing of the plates also is apparent, along with a stair-step topography caused by the stacking of unit layers with different lateral dimensions. These heterogeneous features lead to soil kaolinite aggregates that are not well organized, with many stair-stepped clusters of stacked plates, interspersed with plates in edge-face contact, evidently because of differing surface charge on the edges and faces. Similar observations have been made for 2:1 clay minerals. Illite, for example, has platy particles comprising unit layers stacked irregularly, although the bonding mechanism for the stacking is cross-linking through an inner sphere surface complex of K+, not hydrogen bonding. These particles also exhibit a stair-step surface topography as well as frayed edges produced by weathering. Coatings of Al-hydroxy and humus polymers may be present. Additional complexity comes from nonuniform isomorphic substitutions, with regions of layer charge approaching 2.0 grading to regions with layer charge near 0.5. Smectite and vermiculite have lesser tendency to form colloids comprising extensive stacks because their layer charge is less than that of illite and, therefore, is less conducive to inner sphere surface complexation with K+.

The Soil Solution

The soil solution was introduced in Section 1.2 as a liquid water repository for dissolved solutes. Speaking more precisely, one can define the soil solution as the aqueous liquid phase in soil having a composition influenced by exchanges of matter and energy with soil air, soil minerals, and the soil biota. This more precise concept identifies the soil solution as an open system, and its designation as a phase means two things: (1) that it has uniform macroscopic properties (for example, temperature and composition) and (2) that it can be isolated from the soil profile and investigated experimentally in the laboratory. Uniformity of macroscopic properties obviously cannot be attributed to the entire aqueous phase in a soil profile, but instead is associated with a sufficiently small element of volume in the profile still large enough to include many pores. As is the case for soil humus (Section 3.2), the problem of isolating a sample of the soil solution without artifacts is an ongoing challenge to soil chemistry, but several techniques for removing the aqueous phase from soil into the laboratory have been established as operational compromises between chemical accuracy and analytical convenience. Among these techniques, the most widely applied in situ methods are drainage water collection and vacuum extraction, whereas the common ex situ methods include displacement by another fluid and extraction by vacuum, applied pressure, or centrifugation. The in situ techniques are influenced by whatever disturbance to a soil profile and, therefore, natural aggregate structure and water flow patterns, has occurred because of apparatus installation. They yield a sample of the soil solution from a largely undefined profile volume and they differ in whether they provide the flux compositionor the resident compositionof a soil solution. A flux composition, which is relevant to chemical weathering and, more broadly, to solute transport in soils, is measured in a soil solution sample obtained by natural flow into a collector, as occurs in a pan lysimeter (Fig. 4.1).

Soil Humus

Biomoleculesare compounds synthesized to sustain the life cycles of organisms. In soil humus, they are usually products of litter degradation, root excretion, and microbial metabolism, ranging in molecular structure from simple organic acids to complex biopolymers. Organic acids are among the best-characterized biomolecules. Table 3.1 lists five aliphatic (meaning the C atoms are arranged in open-chain structures) organic acids associated commonly with the soil microbiome. These acids contain the unit R—COOH, where COOH is the carboxyl groupand R represents either H or an organic moiety. The carboxyl group can lose its proton easily within the normal range of soil pH (see the third column of Table 3.1) and so is an example of a Brønsted acid. The released proton, in turn, can attack soil minerals to induce their decomposition (see Eq. 1.2), whereas the carboxylate anion (COO-) can form soluble complexes with metal cations, such as Al3+, that are released by mineral weathering [for example, in Eq. 1.7, rewrite oxalate, C2O42-, as (COO-) 2]. The total concentration of organic acids in the soil solution ranges up to 5 mM. These acids tend to have very short lifetimes because of biocycling, but they abide as a component of soil humus, especially its water-soluble fraction, because they are produced continually by microorganisms and plant roots. Formic acid (methanoic acid), the first entry in Table 3.1, is a monocarboxylic acid produced by bacteria and found in the root exudates of maize. Acetic acid (ethanoic acid) also is produced microbially—especially under anaerobic conditions—and is found in root exudates of grasses and herbs. Formic and acetic acid concentrations in the soil solution range from 2 to 5 mM. Oxalic acid (ethanedioic acid), which is ubiquitous in soils, and tartaric acid (D- 2,3-dihydroxybutanedioic acid) are dicarboxylic acids produced by fungi and excreted by plant roots; their soil solution concentrations range from 0.05 to 1 mM. The tricarboxylic citric acid (2-hydroxypropane- 1,2,3-tricarboxylic acid) is also produced by fungi and excreted by plant roots. Its soil solution concentration is less than 0.05 mM.