Chemical composition of soils

The average chemical composition of igneous rocks is given in Table 11.1. On an oxide basis, silicon and aluminum are first and second in abundance. This reflects the dominance of silicate and aluminosilicate minerals in igneous rocks. Iron is next in abundance in igneous rocks. Next most abundant, are four cations that are important in balancing the charge of silicon-oxygen tetrahedra in minerals-calcium, magnesium, sodium, and potassium. The remaining 1 or 2 percent includes the other elements. The average chemical composition of igneous rocks is similar to the mineralogical composition of many minimally and moderately weathered soils. Such soils contain much quartz, feldspar, and mica in the sand and silt fractions, 2:1 layer silicate clays in the clay fraction, and much of the negative charge of the clays neutralized by the adsorption of calcium, magnesium, sodium, and potassium ions.

ION EXCHANGE

Ion exchange is of great importance in soils. Ion exchange involves cations and anions that are adsorbed from solution onto negatively and positively charged surfaces, respectively. Such ions are readily replaced or exchanged by other ions in the soil solution of similar charge, and thus, are described by the term, ion exchange. Of the two, cation exchange is of greater abundance in soils than anion exchange.

Nature of Cation Exchange

Cation exchange is the interchange between a cation in solution and another cation on the surface of any negatively charged material, such as clay colloid or organic colloid. The negative charge or cation exchange capacity of most soils is dominated by the secondary clay minerals and organic matter (especially the stable or humus portion of SOM). Therefore, cation exchange reactions in soils occur mainly near the surface of clay and humus particles, called micelles. Each micelle may have thousands of negative charges that are neutralized by the adsorbed or exchangeable cations. Assume for purposes of illustration that X- represents a negatively charged exchanger that has adsorbed a sodium ion (Na), producing NaX. When placed in a solution containing KCI, the following cation exchange reaction occurs:

Cation Exchange Capacity of Soils

The cation exchange capacity of soils (CEC) is defined as the sum of positive (+) charges of the adsorbed cations that a soil can adsorb at a specific pH. Each adsorbed K+ contributes one + charge, and each adsorbed Ca 2+ contributes two + charges to the CEC. The CEC is the sum of the + charges of all of the adsorbed cations. (Conversely, the CEC is equivalent to the sum of the - charges of the cation exchange sites.)

Kinds and Amounts of Exchangeable Cations

The major sources of cations are mineral weathering, mineralization of organic matter, and soil amendments, particularly lime and fertilizers. The amounts and kinds of cations adsorbed are the result of the interaction of the concentration of cations in solution and the energy of adsorption of the cations for the exchange surface. The cations compete for adsorption. As the concentration of a cation in the soil solution increases, there is an increased chance for adsorption. The more strongly a cation is attracted to the exchange surface (energy of adsorption); the greater is the chance for adsorption.

The energy of adsorption of specific ions is related to valence and degree of hydration. The energy of adsorption of divalent cations is about twice that of monovalent cations. For cations of equal valence, the cation with the smallest hydrated radius is most strongly adsorbed because it can move closer to the site of charge. The dehydrated and hydrated radi of four monovalent cations are given in Table 11.2. Rubidium (Rb) is the most strongly adsorbed, whereas lithium (Li) is the most weakly adsorbed.

SOIL pH

The pH is the negative logarithm of the hydrogen ion activity. The pH of a soil is one of the most important properties involved in plant growth. There are many soil pH relationships, including those of ion exchange capacity and nutrient availability. For example, iron compounds decrease in solubility with increasing pH, resulting in many instances where a high soil pH (soil alkalinity) causes iron deficiency for plant growth.

Determination of Soil pH

Soil pH is commonly determined by: (1) mixing one part of soil with two parts of distilled water or neutral salt solution, (2) occasional mixing over a period of 30 minutes to allow soil and water to approach an equilibrium condition, and, (3) measuring the pH of the soil-water suspension using a pH meter. A rapid method using pH indicator dye is illustrated in Figure 11.5. The common range of soil pH is 4 to 10.

(H.D. Foth, Fundamentals of Soil Science)