The Chemistry of Soil pH

N. T. Coleman and A. Mehlich.

Every substance dissolved in water or mixed with water is acid, neutral, or alkaline. Vinegar, containing acetic acid, is a substance that is acid. Water itself is neutral, as are solutions of salt and other such compounds. Lime, baking soda, and lye give alkaline solutions.

The reactions of these and all other substances are typical of the substances themselves and of the ways in which they react with water.

Water itself is composed of one atom of oxygen and two atoms of hydrogen. Its chemical formula is H₂O. Liquid water, however, is really a mixture of H₂O, H+, and OH-. H+, called the hydrogen ion, is a positively charged hydrogen atom, and OH-, the hydroxyl ion, is a negatively charged unit consisting of one hydrogen and one oxygen atom.

The hydrogen ion and hydroxyl ions come from the breakdown, or ionization, of the water molecule according to the scheme:

H₂O<->H++OH-.

This reaction does not produce very many H+ or OH- ions because the water molecule is very stable. In fact, only about one molecule in 10 million is ionized at any one time.

The extent to which water ionizes can be expressed more exactly in terms of an ionization constant, Kw. This is defined as:

Kw=[H+] [OH-]. (I)

[H+] and [OH-] are the concentrations of hydrogen ions and hydroxyl ions, expressed as equivalents per liter. One equivalent of a singly charged ionic species is the weight in grams of that species, which contains 6.023 x 10²³ particles.

The value of Kw is 10^-14 at 22 C., and in any aqueous system at this temperature the product of the concentrations of hydrogen and hydroxyl ions is 10^-14.

Equation I can be written in a more useful form by dividing both sides into I and taking logarithms. That is:

The values of log I /[H+] and log I [OH-] generally are referred to as pH and pOH, respectively. These values are indices of the acidity or alkalinity of a system.

Any system in which pH and pOH are equal is neutral. At 22 C., when Kw=10-14, neutrality corresponds to pH=pOH=7. When pH is less than 7, the system is acid. When pH is above 7, it is alkaline.

Soils vary in pH from about 4, for strongly acid soils, to about 10, for alkaline soils that contain free sodium carbonate.

The pH range for most agricultural soils is 5 to 8.5.

IN SOLUTIONS, pH is related to hydrogen-ion concentration in a straightforward manner. That is not the case for soils, which consist of a solution phase, the soil water, and a solid phase, the mineral and organic particles of the soil. The pH of a soil-water system is an approximate reflection of the hydrogen-ion concentration of the soil solution, but it does not reflect the total acidity of the system. This is because of the cation-exchange properties of soils.

The tiny mineral and organic particles of soils have cation-exchange capacities the particle surfaces are negatively charged, and the positively charged ions, or cations, sit on or near the particle surfaces. These positive ions are called exchangeable cations.

The cation-exchange capacity of a soil is the quantity of positive ions necessary to neutralize the negative charge of a unit quantity of soil, under a given set of conditions. Cation-exchange capacity usually is expressed as milliequivalents (6.023 x 10²⁰ particles) of cations required to neutralize the negative charge of 100 grams of soil at PH 7.

The cation-exchange capacity of a soil depends on the amounts and kinds of finely divided mineral and organic particles present. Sandy soils generally have low cation-exchange capacities because of their small proportions of negatively charged material. Soils high in organic matter have substantial cation-exchange capacities because of the large negative charge developed by humus. As far as clays are concerned the cation-exchange capacities of the montmorillonoid and vermiculite-like minerals, found in Midwestern soils and soils of the dry areas, are large. Those of kaolin minerals, which predominate in Southeastern soils, are small. The finely divided micas are intermediate.

The fine-grained mineral and organic particles of soil differ as to the source and characteristics of their negative charge as well as to its magnitude. There are two general sources of cation-exchange capacity, a permanent charge and a pH-dependent charge.

All clay minerals appear to possess permanent charges resulting from their crystal structures. As the name implies, the permanent charge of soil minerals persists under all conditions. The montmorillonoid minerals, which occur in many soils of the Midwest, have permanent charges of 80 to 120 milliequivalents per 100 grams. The kaolin minerals, which are predominant in soils of the Southeastern States, have much smaller permanent charges, varying from 1 to about 8 milliequivalents per 100 grams.

Clay minerals also have pH-dependent charges, which result from the ionization of hydrogen ions from SiOH groups around the edges of the crystals and perhaps for other reasons. Such charges do not develop below a PH of about 6, and do not contribute to the effective cation-exchange capacity of acid soils. As pH is increased above 6, the pH-dependent charge increases progressively, reaching a maximum at a pH of around 10.

The pH-dependent charge which can be developed by clays depends on the exposed edge area of the crystals and on the nature of the clay. It probably is around 20 milliequivalents per 100 grams for montmorillonoid clays, and is much smaller for micas and kaolins. Cation-exchange capacity at PH 7 includes all of the permanent charge and a part of the pH-dependent charge.

Because of the chemical nature of soil organic matter, it probably has only pH-dependent charge, resulting from the ionization of hydrogen from carboxyl and phenol groups attached to organic matter particles. Of these two kinds of groups, the carboxyls ionize largely between pH 4 and pH 7, while the phenol groups ionize only at alkaline reactions. The cation-exchange capacity of soil organic matter, measured at pH 7, averages near 200 milliequivalents per 100 grams. Thus I percent of organic matter contributes about 2 milliequivalents per 100 grams toward the cation-exchange capacity of a soil.

The relationships between cation-exchange capacity and pH for clay minerals and organic matter can be illustrated schematically, as in figure I.

For the mineralogical clay specimens the effective cation-exchange capacity does not vary with pH below about 5. This may not be the case with soil clays. The effective cation-exchange capacity of organic matter is near zero at a pH below about 4, and increases continuously as the pH is raised.

Though clay minerals and soil organic matter generally are considered to be responsible for the cation-exchange capacity of soils, other substances may contribute as well. Iron, aluminum, and titanium oxides, as well as noncrystalline iron and aluminum silicates and phosphates, may add greatly to the cation-exchange capacities of some soils.

1. The negative charge, or effective cation-exchange capacity, of clays and organic matter varies with pH. Those of kaolinite (curve 1) and montmorillonite (curve 2) are constant below pH 6 but increase at more alkaline reactions. The negative charge of humus (curve 3) increases linearly with pH.

Under any given set of soil conditions, the permanent charge and that part of the pH-dependent charge that is ionized are balanced by the presence of cations in the vicinity of the particle surfaces. The kinds of balancing ions, or exchangeable cations, depend on the previous history of the soil.