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Chapter 7 Soil minerals and weathering.

"Probably more harm has been done to the science by the almost universal attempts to look upon the soil merely as a producer of crops rather than as a natural body worth in and for itself of all the study that can be devoted to it, than most men realize."

C. F. Marbut, 1920

Soil chemistry is
important because it
affects fertility, leaching,
waste remediation, and
shrink-swell potential.

You have finished reading about many important physical properties of soil and the physical interactions among soil, water, and air. However, it is the chemical properties of soil that greatly determine the role of soil in plant growth, waste treatment, and other important areas. The minerals in the soil and their associated chemical properties control the ability of soil to adsorb and retain ions and water. Soil chemistry influences soil fertility and plant nutrition by controlling nutrient-holding capacity and leaching potential. Soil chemistry influences the capacity of soil to remediate waste effluent. Soil chemistry influences soil physical properties such as soil shrinkage, swelling, and cohesiveness. Soil chemistry influences the fate of chemicals added to soil, and how much and by what methods amendments should be applied.

To introduce the chemistry of the soil environment, in this chapter you will examine the types of minerals found in soils and how those soil minerals develop through weathering. You will examine the chemical properties of these minerals and how these mineral properties, as well as soil management, influence the chemical properties of the soil solution.


After reading this chapter, you should be able to:

* List five ways that soil chemistry influences soil properties.

* Understand the difference between primary and secondary minerals.

* Identify key chemical properties of soil minerals.

* Know four important outcomes of weathering in soil.

* Identify several types of colloids.

* Describe two ways by which soil colloids develop charge.


base cations





isomorphous substitution



primary mineral

secondary mineral


specific surface area




The mineralogic
composition of parent
material determines the
mineralogic composition
of the resulting soil.

Soils form from rocks, and rocks are composed of minerals. The chemical properties of a soil can therefore be traced back to the chemical properties of the minerals within the original rocks. This can be seen if the original minerals still persist in soil and if weathering has transformed the primary minerals into secondary minerals. The nature of the primary minerals influences the type of secondary minerals that may develop. In many cases, soil chemical properties such as nutrient availability and pH can be predicted from the parent material mineral properties. Physical properties such as texture and degree of development may also be predicted from parent material mineralogy. You will begin this study by briefly examining rock and mineral types and weathering processes before you turn your attention to the chemical properties of the soils that result from this weathering.

Mineral Composition

Approximately 75 percent of the weight and approximately 95 percent of the volume of the Earth's crust (upper 10 miles) consists of oxygen (O) and silicon (Si; Table 7-1). Consequently, the types of minerals that dominate the rocks and soils of the Earth's surface are silicates and aluminosilicates (Table 7-2).

As was discussed in Chapter 4, there are differences in the mineral composition among the different-sized soil particle separates. Sand and silt particles predominantly contain primary minerals. Sand particles are mostly quartz grains, which are highly resistant to weathering. Silt particles are also typically grains of quartz, with other resistant primary silicate minerals (such as orthoclase feldspar) also being common. Clay particles, however, are dominated by secondary minerals. These develop from the weathering of the less resistant primary minerals (such as anorthite feldspar or olivine).


Weathering is a complex process by which rock is transformed into regolith and regolith is transformed into soil. Physical weathering transforms larger particles into smaller particles. Chemical weathering transforms primary minerals into secondary minerals.

Physical Weathering

Physical weathering is characterized by the mechanical disintegration of rocks. Physical weathering alone does not change the chemical nature of the minerals in the weathered materials. Physical weathering occurs when pressure or stress is applied to the rock or other mineral material and then released. The source of this stress may be from an external agent, such as ice. During freezing and thawing cycles, for example, the expansion of ice that forms within small cracks in a rock may exert a pressure of 150 T/[ft.sup.2] (1465 Mg/[m.sup.2]) as it expands. Expanding roots may also exert physical stress on cracks in rocks. The stress may come from within the rock as it expands and contracts with heating and cooling.

Abrasion by water, ice, or sediment-carrying wind can contribute to the physical disintegration of rocks. Rocks that have been buried under large masses of material (including other rock layers) will expand following the release of the overburden stress. Chemical weathering may also change the volume of a mineral and put stress on the mineral structure, leading to physical weathering. In general, physical weathering decreases the size of mineral particles and increases their surface area. The increased surface area makes the minerals more susceptible to chemical weathering reactions, which mainly occur at mineral surfaces.

Chemical Weathering

Chemical weathering is the decomposition of minerals in response to [H.sub.2]O, [O.sub.2], C[O.sub.2], and other chemicals. During the weathering process the elemental composition of the rocks or soil changes (Table 7-3). Initially, the basic elements (K, Na, Mg, Ca) are readily lost, along with some loss of Al and Fe. Over time, though, as weathering proceeds to greater extents, Si is also lost. In intensively weathered soils, Al and Fe have become concentrated in the soil while other elements have been lost.

Common weathering processes include hydration, dissolution, and hydrolysis, all of which are driven by the presence of water. Hydration occurs when one or more water molecules are attached to the molecular structure of a mineral. A simple hydration reaction is the addition of water to hematite:

[Fe.sub.2][O.sub.3(s)] + 3[H.sub.2][O.sub.(1)] [right arrow] 2Fe[(OH).sub.3(s)]

This reaction changes the hematite from a red color to a yellowish-brown color.

A more important result of some hydration reactions is that adding water changes the volume of the mineral. As mentioned before, this puts stress on the mineral structure, leading to physical weathering or to further chemical weathering. Among common minerals in the soil, gypsum is normally found as a hydrated mineral:

CaS[O.sub.4(s) + 2[H.sub.2][O.sub.(1)] [right arrow] CaS[O.sub.4] x 2[H.sub.2][O.sub.(s)]

With further addition of water, the gypsum will undergo dissolution.


Mineral dissolution occurs when the solid materials dissolve in water. This is commonly observed when salt dissolves in water:

[NaCl.sub.(s)] + [H.sub.2][O.sub.(l)] [right arrow] [Na.sup.+].sub.(aq)] + [Cl.sup.- ].sub.(aq)] + [H.sub.2][O.sub.(1)]

While not shown explicitly, in this and similar reactions hydration may be an important precursor to dissolution. The addition of the water molecules weakens the bonds between the elements within the mineral.

Hydrolysis involves splitting water into [H.sup.+] and O[H.sup.-]:

[H.sub.2][O.sub.(l)] [right arrow] [H.sup.+.sub.(aq)] + O[H.sup.-.sub.(aq)]

This reaction sometimes follows a hydration reaction. For example, when aluminum is released from a mineral it readily undergoes hydrolysis:


This reaction will be given more attention later when you examine soil acidity.

In terms of common soil minerals, calcite weathering provides another example of hydrolysis:


This reaction, though, is a simplified view of a more complex series of reactions that likely occur during weathering, including hydrolysis, dissolution, and hydration. Particularly when hydrolysis is involved, there is an exchange of hydrogen for other cations in the mineral structure (for example, [H.sup.+] for [Ca.sup.2+] in calcite). Consequently, the presence of [H.sup.+] tends to accelerate weathering.

In the more complex soil system, weathering is commonly through the attack of [H.sup.+] on the A1-O bonds in clay minerals:


Here, the weathering of a primary mineral, albite, is driven by the presence of [H.sup.+] in an aqueous system. Further hydrolysis of water alters the albite to a secondary mineral, kaolinite, and releases the sodium and some soluble silica.

There are at least four important products of the combined weathering reactions in the soil. The results of weathering reactions include: (1) formation of clay-sized secondary mineral particles [for example, [A1.sub.2][Si.sub.2][O.sub.5][(OH).sub.4]], (2) the release of ions into the soil solution (for example, [Na.sup.+]), (3) the formation of other soluble compounds that may be leached from the soil [for example, Si[(O[H).sub.4]], and (4) the concentration of residual resistant minerals.

Factors Affecting Weathering Rates

The factor that most influences weathering rates is climate, including temperature and precipitation. With increasing temperatures there is an increase in chemical reaction rates, so at higher temperatures weathering reactions are more rapid. Similarly, greater precipitation and greater moisture content increase weathering rates. Water is an important reactant in almost all chemical weathering reactions. In arid environments there is little weathering beyond physical weathering.

The physical and chemical characteristics of the initial minerals are also important. Some minerals are more resistant to weathering than others (Table 7-3). This resistance is derived from the strength of bonding in the mineral structure and the particle size of the mineral grains in the material. For sedimentary rocks, the degree of cementation is also important.

There is also a biological contribution to weathering. Carbon dioxide dissolves in water, becomes hydrated, then hydrolyzes the water:


Root and soil organism respiration provides a ready source of C[O.sub.2] in soil environments for this reaction to occur, and lowers the pH of the soil solution. Similarly, the presence of organic acids from organic matter decomposition increases acidity and promotes weathering. Plant nutrient uptake may also change the chemical equilibrium in soil and promote further mineral decomposition reactions that release additional nutrient ions.


1. What is weathering? What are the two general types of weathering? What is the primary difference between the two?

2. What is the difference between primary minerals and secondary minerals?

3. Do all chemical weathering reactions involve water in some way? In regard to the water, what is the difference between hydration and hydrolysis? Between these and dissolution?

4. Why would chemical weathering proceed much more slowly without the living organisms on Earth?

5. What new substances may be formed from those released during chemical weathering?


Clay formation is an important outcome of weathering reactions. You have already seen the term clay used to refer to a soil particle size separate and a soil textural class. You will also often use the term clay to refer to a class of soil minerals, namely the layer silicate clays. Chemical weathering processes alter primary minerals into these clay-sized secondary minerals. These particles impart many important properties to the soil.

Clays and Colloids
Many fundamental soil
processes are surface

The defining property of a clay particle, as discussed in Chapter 4, is size. By definition, clay particles are small, with effective diameters less than 0.002 mm. The smallest of the clay particles are termed colloids. These particles are less than 0.001 mm in diameter. In addition to their small size, clay particles differ in shape from the larger sand and silt particles. Clay particles have a platelike shape (Figure 7-1), which is due to the chemical structure of the minerals. Individual clay particles will often become oriented in the soil, forming stacks, or domains (Figure 7-2).


Many clay particles
develop surface charge
because of changes in
chemical structure.

The small size and flat shape of the clay particles imparts an extremely large specific surface area (surface area per gram of soil) to these particles. Because the clay particles are often found in domains, the surface area can be divided between internal surface area (between individual particles) and the external surface area (on the exposed edges of the individual particles; as in Figure 7-2). The vast majority of the surface area of clay particles is internal surface area.

Layer silicate clay minerals also have important chemical properties. Due to an imbalance of charge in the chemical structure of the minerals, many clay particles develop surface charges. These charges are predominantly negative, but under some circumstances a positive charge can develop. These charges are important because they result in a soil's ability to adsorb positive ions (cations such as [H.sup.+], [K.sup.+], [Ca.sup.2+], [Mg.sup.2+]). Polar water molecules are also attracted to the negative charges on the clay and colloid surfaces. As a result, the clay and colloid particles are important components of the soil that lead to cation and anion exchange, promote soil structure formation and stability, and influence water retention and movement.

Types of Colloids
Colloids can be mineral or
organic and crystalline or

There are actually several types of mineral and organic soil colloids. The most common in most soils are the layer silicate clays. The properties of these crystalline clay minerals are quite variable. Consequently, clay type is similar to clay content in its influence on soil physical and chemical properties. All are characterized by the presence of an ordered crystalline structure, most commonly forming thin mineral layers that stack upon one another. You will get more details on clay mineral types in later sections.
Iron and aluminum oxides
are important soil

A second general type of mineral colloid is referred to as amorphous (noncrystalline) minerals (for example, imogolite and allophane). These minerals lack the ordered crystalline structure found in the layer silicate clays. They often have high surface charges, which may be negative or positive, giving them the ability to adsorb nutrients and water. These minerals, though, are not common in most soils. They are most often associated with soils formed from volcanic parent materials.

Other types of mineral colloids are the oxides of iron and aluminum (for example, hematite, goethite, and gibbsite). These minerals are most common in highly weathered soils, such as those found in the southeastern United States and tropical regions of South America, Africa, and southeast Asia. Unlike other soil colloids, these minerals have low surface charges and therefore low ability to adsorb nutrients and other cations. The iron oxides in particular are important because they are the source of red, orange, and yellow colors in most soils, with these colors becoming more pronounced as the amount of iron oxides increases.
Hydroxyl and carboxyl
groups are important
sources of charge on

A fourth important soil colloid type is the organic soil colloids, or humus. These colloids are found in varying amounts in almost all soils--the product of plant and animal decay in or on soil. Humus is noncrystalline, but has a high surface charge that is predominantly negative. The source of this charge is the hydroxyl (--OH), carboxyl (--COOH), and similar functional groups that are components of these large, complex organic molecules. As you will see in later discussion, the charge on these organic colloids highly depends on the soil pH.


1. What is a colloid?

2. What properties distinguish colloids from other soil particles?

3. In reference to soil silicate clay particles, what is the difference between internal and external surface area?

4. What are the four major types of colloids present in soils? What are some important properties of each?

5. In what areas/environments is each type of colloid most commonly found?

6. What are common cations that are adsorbed to colloids? In what areas/environments is each most commonly found?


Now it's time to turn your attention to the dominant type of colloid found in most soils. The characteristic physical and chemical properties of the layer silicate clays are directly related to the chemical structure of these minerals. You will examine the basic structure of these minerals, which will then explain the source of their charge and the cause of the stickiness, plasticity, and swelling capacity commonly associated with clay.

Layer Silicate Clay Structure
Layer silicate clays consist
of sheets of Si, Al, Mg, Fe,
and other elements held
together by oxygen atoms
and hydroxyl groups.

Layer silicate clays are ordered crystalline minerals, consisting mostly of layered sheets of Si, Al, Mg, and/or Fe atoms surrounded and held together with oxygen atoms and hydroxyl groups. The differences among the various clay mineral types are dictated by the composition and arrangement of these sheets.

Tetrahedral and Octahedral Sheets

Two types of crystalline sheets may be found in the crystalline structure of a layer silicate clay. The silicon tetrahedral sheet has a silicon atom bonded to four oxygen atoms as its primary unit (Figure 7-3). The four oxygen atoms, seeking maximum separation from each of the other oxygen atoms, can be seen as the points of a four (tetra-) sided (-hedron) solid (Figure 7-4). The tetrahedron is sometimes described as a three-sided pyramid. A silicon tetrahedral sheet is formed when a Si atom shares each of its oxygen atoms that make up the base of the pyramid (basal oxygen atoms) with three adjacent Si atoms (Figure 7-5).






The Al-Mg octahedral sheet has an Al or Mg atom bonded to six oxygen atoms as its primary unit (Figure 7-6). In this case, the six oxygen atoms can be seen as the points of an eight (octa-) sided (-hedron) solid (Figure 7-7). This is sometimes described as two four-sided pyramids joined at their bases. An octahedral sheet is formed when an Al or Mg atom shares four or six of its oxygen atoms with three or four adjacent Al or Mg atoms (Figure 7-8).

Two or more sheets of
layer silicate crystals
joined together are called
a layer.

Soil clay minerals are not made up of individual tetrahedral or octahedral sheets. Layer silicate clay crystals consist of two or three of these sheets joined together (through the sharing of oxygen atoms) in one of several specific arrangements. When two or more sheets are joined together the resulting unit is termed a layer. The most simple configuration is for one octahedral sheet to be joined to one tetrahedral sheet. This occurs through the sharing of the apical oxygen atoms in the tetrahedral sheet.


Another common configuration is for one octahedral sheet to be joined to two tetrahedral sheets. The octahedral sheet is sandwiched between the two tetrahedral sheets, again joined through the sharing of the apical oxygen atoms of the tetrahedral sheets. For both of these layer combinations it is important to note that the structural arrangement of the atoms is very precise. However, the atoms that occupy these arrangements (not including the oxygen atoms) can change in response to various factors during the crystallization, alteration, weathering, and recrystallization of these minerals.


1. What is a tetrahedron? What atoms form the silicon tetrahedra in a layer silicate clay crystal?

2. What is an octahedron? What atoms form the aluminum-magnesium octahedra in a layer silicate clay crystal?

3. Tetrahedra or octahedra are joined together in sheets (tetrahedral or octahedral). When two or more sheets are stacked together, what is this called?


Under ideal circumstances, both the individual sheets and the combined layers of the layer silicate clay minerals have no net charge, because the negative charges of the oxygen atoms are exactly balanced by the Si, Al, Mg, and hydrogen atoms in the chemical structure of the mineral. In reality, as these minerals form, the makeup of the crystals may diverge from the ideal. The crystals are also influenced by the surrounding conditions (mainly pH), which may change the structural makeup. In both of these cases the number of oxygen atoms remains unchanged. Adding or replacing the other atoms alters the number of positive charges in the crystal. With the number of negative charges from the oxygen atoms unchanged, the result is a change in the net charge on the crystal. These sources of charge on clay minerals and organic colloids lead to the net negative charge that attracts cations and water to these colloids.

Constant Charge
Isomorphous substitution
occurs when atoms of
similar size but different
charge replace elements in
silicate clay layers.

In layer silicate clay minerals, the primary source of charge arises when the Si, Al, or Mg atoms in the crystal are replaced by atoms that are approximately the same size, but of different charge. When this happens, the balance of positive and negative charges changes, but the arrangement of atoms in the crystal does not. For this reason, the process is called isomorphous substitution (iso = same; morpho = shape). For example, replacing a silicon atom ([Si.sup.4+]) with an aluminum atom ([A1.sup.3+]) in the tetrahedral sheet (Figure 7-9) produces a net charge of - 1. Conversely, replacing a magnesium atom ([Mg.sup.2+]) with an aluminum atom ([A1.sup.3+]) in the octahedral sheet (Figure 7-9) produces a net charge of + 1. In both of these cases, these substitutions produce permanent charges within the clay mineral structure. However, not all substitutions affect the net charge. Replacing an aluminum atom ([A1.sup.3+]) with an iron atom ([Fe.sup.3+]) in the octahedral sheet (Figure 7-9) produces no net charge.


Variable Charge
Variable charge depends
on the pH of the soil
when acid and negative
when basic.

Hydroxyl (-OH) groups are found throughout the crystal structure of layer silicate clay minerals. Similarly, humus molecules contain many hydroxyl groups. As was seen with the water molecule (Chapter 6), the strong electronegativity of the oxygen atom leads to unequal sharing of the electron in the covalent bond between hydrogen and oxygen in the hydroxyl group. The electron of the hydrogen atom is drawn relatively closer to the oxygen atom (polar covalent bonds). This creates a negative charge concentration near the oxygen atom in the hydroxyl group. Because of this polarity, the hydroxyl groups of the colloid can interact with hydrogen ions or hydroxyl ions in the soil solution.

At low pH, when the concentration of H+ ions in the soil solution is high, the electronegative oxygen atoms in the colloid can attract an electropositive H+ ion in the soil solution (Figure 7-10). By adding a + 1 hydrogen ion to the composition of the colloid, the net charge becomes more positive (less negative). At high pH, when the concentration of O[H.sup.-] ions in the soil solution is high, the electronegative oxygen atoms of the O[H.sup.-] ion can attract an electropositive [H.sup.+] ion in the colloid (Figure 7-10). By removing a + 1 [H.sup.+] ion from the composition of the colloid the net charge becomes more negative (less positive).
Most of the permanent
charge of clay minerals
comes from isomorphous

The various soil colloids have differing amounts of total charge, as well as differing amounts of constant versus variable charge. Organic colloids have the greatest charge per gram of colloid, with most of this charge being pH-dependent. Iron and aluminum oxides also have predominantly variable charge, but their total charge is quite low. The layer silicate clays, in general, derive most of their charge as permanent charge through isomorphous substitution. However, each of the clay mineral types varies in their total charge and in their charge distribution between constant and variable forms. For this reason it is important to appreciate the different types of clay minerals. They are just as important as the total amount of clay in determining the chemical properties of a soil.



1. What are the two major sources of charge on soil colloids? What is the fundamental difference between them?

2. What is isomorphous substitution? How is it related to the charge of soil clay particles?

3. If there is a substitution of [Mg.sup.2+] for [Al.sup.3+] in an octahedral sheet of a layer silicate clay, does the net negative charge of the clay increase or decrease?

4. What is more common among layer silicate clays, a net negative or net positive charge?

5. What is the source of pH-dependent, or variable, charge on soil colloids?

6. Chemically, what happens to the hydroxyl (-OH) group under basic (pH > 7) conditions? How does this change the net charge on the soil colloid?

There are several different
types of silicate clay
minerals: illite, vermiculite,
smectite, and kaolinite.

There are two general types of silicate clays. Minerals with one tetrahedral sheet and one octahedral sheet are termed 1:1 clays. Minerals with two tetrahedral sheets for each octahedral sheet are termed 2:1 clays. All are considered secondary minerals, formed by weathering of primary minerals. In fact, there are general geologic and climatic conditions that promote formation of the different clay minerals. Accordingly, each of the different silicate clay minerals has different elemental compositions. We have noted that some of the important properties of clay particle include relatively high surface area, the presence of negative surface charges, stickiness, and the ability to swell upon wetting. Each of the common clay minerals differs in these key properties. We consider here the four most common types of soil clay minerals.


Illite (Figure 7-11) is a 2:1 type clay with moderate surface area (100-120 [m.sup.2]/g) and moderately high negative surface charge. The charge arises through isomorphous substitution of [A1.sup.3+] for [Si.sup.4+] in the tetrahedral sheets. This concentration of charge on the outer surfaces of the individual clay layers promotes the strong adsorption of cations, mainly potassium ([K.sup.+]). The presence of the [K.sup.+] creates strong bonds between individual clay layers, causing illite to be only slightly sticky with low to moderate swelling potential.

The illite clay minerals are normally associated with less intense weathering environments where the parent materials were relatively high in base cations ([Ca.sup.2+], [Mg.sup.2+], [K.sup.+], and [Na.sup.+]) and climatic conditions have not favored the leaching of these base cations.



Vermiculite (Figure 7-11) is also a 2:1 type clay, but with a high surface area (700-800 [m.sup.2]/g) and a high negative surface charge (100-200 meq/100 g soil or 100-200 [cmol.sub.+]/kg soil). Similar to illite, this charge originates mainly from isomorphous substitution in the tetrahedral sheets, although some substitution in the octahedral sheet does occur. Though not as prominently as with illite, this charge configuration promotes stronger adsorption of interlayer cations. There exist moderately strong bonds between the individual crystal layers, yielding a sticky consistency and a moderate swelling potential for vermiculite clays. Vermiculite minerals are also associated with less intense weathering environments, but have begun with or lost more of the potassium, which helps hold the clay layers together in illite.

Smectite (Montmorillonite)

Smectite (Figure 7-11) is the third prominent 2:1 type clay found in many soils. Similar to vermiculite, it has a very high surface area (especially when wet; 600-800 [m.sup.2]/g) and a high negative surface charge (80-150 meq/100 g soil; 80-150 [cmol.sub.+]/kg soil). Unlike illite and vermiculite, this charge originates from isomorphous substitution in the inner octahedral sheet. Consequently, the interlayer cations that can help bond clay layers together are not held as tightly as with illite and vermiculite.

Smectite clays have weak bonds between layers, which gives them a very sticky consistency and a high swelling potential. This propensity to swell upon wetting makes management of soils with high smectite clay contents a significant challenge. Construction activities are particularly hampered by the swelling and shrinking cycles that occur, which can damage foundations, crack walls, shift poles, and warp roadbeds.

Smectite minerals are more associated with moderately weathered environments where there has been greater removal of base cations over time, particularly K and Mg.


Kaolinite (Figure 7-11) is a 1:1 type clay with low surface area (5-30 [m.sup.2] per gram of soil) and low negative surface charge (1-15 milliequivalents per l00 grams of soil; 1-15 [cmol.sub.+]/kg soil). The low charge is a result of limited isomorphous substitution in the crystal structure. When stacks of individual kaolinite crystal layers occur, the oxygen atoms of the exposed tetrahedral sheet are adjacent to the hydrogen atoms of the hydroxyl groups of an octahedral sheet. As a result, relatively strong hydrogen bonds develop between the layers, causing kaolinite to have a very low swelling potential (almost none) and to be only slightly sticky.

Soil kaolinite is mainly found in leached environments with acid conditions associated with hot, wet climates. Consequently, it is the most common clay mineral in the southeastern United States.

The nonexpansive nature of kaolinite clay contributes to its many and varied uses as an extractable resource. It has long been used for making bricks, tiles, and other earthenware, from simple pottery to expensive china and ceramics. Kaolinite can also be used to produce smooth, shiny coatings on glossy paper and soothing coatings on your stomach to combat diarrhea, and is sometimes used in paints as a pigment and in nondairy products (shakes and ice cream) as a thickening agent.

Nonsilicate Clays
Nonsilicate minerals don't
have isomorphous
substitution but they do
have pH-dependent

The various Fe and Al oxide minerals, including hematite, goethite, and gibbsite, are the most common nonsilicate clay colloids found in most soils. There is no isomorphous substitution in these minerals, so they possess very low negative surface charge (0-15 meq/100 g soil; 0-15 [cmol.sub.+]/kg soil), and this charge is entirely pH-dependent. At low pH, the oxygen atoms may readily adsorb hydrogen ions from solution and a net positive charge may develop. These minerals lack the defined crystal structure associated with the layer silicate clays so they are nonsticky and have no swelling potential. These minerals are the result of highly intense weathering environments where most if not all of the base cations and much of the silicon has been lost. The Fe oxide minerals are noteworthy because they are a primary source of color in most soils. This explains the typical strong red colors associated with the highly acid, highly leached soils common in the humid tropical regions, as well as the general lack of nutrient-holding capacity and phosphorus availability.

Clay Minerals and Their Distribution

Soil mineralogy is not the same everywhere. It is affected by the distribution of soil parent materials, from which the soil minerals originate, and it is affected by the distribution of soil-forming processes that lead to the deposition, alteration, decomposition, formation, and movement of soil particles. Differing mineralogy with depth and from one area to another can influence land use because it affects:

* water relations, particularly moisture availability

* physical properties, such as consistency (stickiness, plasticity) and expansiveness

* chemical relations, such as nutrient storage, buffering capacity, surface activity (adsorption potential)

* solute transport and leaching potential

Distribution and Landscapes

The assortment of minerals in a soil changes with time in response to the intensity of chemical weathering. Although conditioned by the mineralogy of the soil parent materials, some general trends exist (Table 7-2; Table 7-4). The most soluble minerals, gypsum and carbonates (calcite, dolomite), are normally only found in young sediments or in arid soil environments where moisture is lacking. Among the clay minerals, you have already seen the general weathering progression from illite and vermiculite, to smectite, to kaolinite, to Fe and Al oxides.

As weathering intensity increases, the layer silicate clays are removed and the relative abundance of Fe and Al oxides increases. Quartz, a rather resistant primary mineral, is relatively abundant in mildly weathered soils. Its relative abundance increases as less-resistant minerals are weathered and removed from the soil. However, over time, even quartz begins to be weathered away. These general trends in weathering sequences and products let you make inferences about soil-forming environments and weathering intensity based on the assortment (and amounts) of various minerals in a soil.

Water moving over and through the soil can produce differences in the local weathering environment and yield differences in soil mineralogy along hill slopes and across landscapes. There are no general rules that can be applied to all landscapes, but patterns do exist and can be recognized and explained. In upper landscape positions the leaching intensity is greater, so weathering products (bases, soluble silica) are easily lost. These products may then accumulate in lower landscape positions, where leaching is less intense. As a result, the upper slope positions tend to have more of the clay minerals associated with intensive weathering (Al oxides, kaolinite) and the lower horizons may have more of the clay minerals associated with less intensive weathering (smectite, vermiculite).

Distribution and Depth
Weathering reduces the
amount of silicate clay
minerals and increases
the amount of iron and
aluminum oxides.

The intensity of weathering reactions and the results of subsequent soil-forming processes may also alter the distribution of soil minerals with depth. In intermediate to strongly weathered soils in humid climates, the formation of secondary clay minerals and subsequent downward translocation of clay not only alters soil texture (as discussed in Chapter 4), but will also concentrate primary minerals in the upper horizons (A and E) and enrich the subsoil horizons in secondary clay minerals (Table 7-5).

Variation in the specific types of clay minerals that occur at different depths within the soil profile is controlled by the leaching intensity. In arid areas with little leaching, or in high-precipitation areas with excessive leaching, clay type is more or less uniform throughout the profile. In a more moderate leaching environment the weathering products from the upper horizons are translocated to the subsoil. As a result, upper horizons may have more of the clay minerals associated with intensive weathering (Al oxides, kaolinite) and the lower horizons where base cations and silicates accumulate may have more of the clay minerals associated with less-intensive weathering (smectite, vermiculite).


1. What is the difference between a 1:1 and a 2:1 layer silicate?

2. How are the sheets of a layer silicate held together? How are adjacent layers of a clay crystal (or micelle) held together?

3. What causes clay layers to expand when wetted? Why are micas nonexpanding?

4. Rank the following clay minerals in terms of their net charge: vermiculite, smectite, mica.

5. Which clay minerals are most common in soils with low weathering intensity?

6. Which clay minerals are most common in highly weathered soils?


Soil chemistry begins with the minerals of soil and how they form through weathering, which influences soil fertility and plant nutrition by controlling nutrient-holding capacity and soil physical properties such as soil shrinkage, swelling, and cohesiveness. Primary minerals in sedimentary, metamorphic, and igneous parent materials weather by physical and chemical processes to yield secondary minerals. Secondary minerals have smaller size and greater surface area, which increases their chemical reactivity.

The secondary minerals produced by weathering are composed of layer silicate materials or amorphous Fe and Al oxides. The layer silicates are composed of oxygen, Si, Al, and other elements arranged into tetrahedral or octahedral sheets. These sheets are organized into layers, and the layers are organized into stacks or domains.

Layer silicates can have charge because of isomorphous substitution in which atoms of similar size, but lesser charge, substitute for elements within the layer silicates. This is the most important source of permanent charge in clay minerals. There is also variable pH-dependent charge in organic and mineral colloids that is primarily due to hydroxyl groups.

Weathering forms secondary minerals and soil colloids, but as weathering proceeds the secondary minerals begin to be translocated from the soil surface into the soil profile. As weathering proceeds, even Si in the clay minerals is leached and most of the colloids in soil are composed of Fe and Al oxides, which have low cation and anion exchange capacity and therefore reduced nutrient-holding capacity. Illite, vermiculite, smectite, and kaolinite are four significant types of 1:1 and 2:1 clay minerals in soil.


1. Why is physical weathering more prevalent in dry, cool regions, while chemical weathering is more prevalent in wet, hot regions?

2. How does temperature contribute to physical weathering? How does abrasion contribute to physical weathering? How do plants and animals contribute to physical weathering?

3. Which type of soil would be better for growing plants: an immature soil high in primary minerals, or a mature soil high in secondary minerals? Explain your answer.

4. Why do you think that [A1.sup.3+] is held more tightly by soil colloids than [Na.sup.+]?

5. Why does the relative concentration of cations in the soil solution influence the proportion of cations adsorbed to colloids?

6. Which is better for building roads or foundations, kaolinite or smectite? Explain your answer.

7. How is the relative amount of interlayer expansion related to the net charge of the clay layers?

8. How do a soil's local climate and parent materials influence the weathering stage of its clay minerals and, therefore, the geographic distribution of clay minerals?

9. Most soils exhibit a net negative charge. Under what conditions might a soil exhibit a net positive charge?

10. Consider a clay mineral with the following structural formula: ([Al.sub.1.5][Mg.sub.0.5])([Si.sub.3.7][Al.sub.0.3])[O.sub.10](O[H).sub.2]. Which are the tetrahedral atoms? Which are the octahedral atoms? What is the net charge on this clay mineral?

11. Consider a soil with 4 percent organic matter and 20 percent clay (5 percent smectite and 15 percent kaolinite). Using the representative charges shown in Table 8-1, calculate an estimate of the cation exchange capacity (at pH 7) of this soil.

12. Chemically, what happens to the hydroxyl (-OH) group under acidic (pH < 7) conditions? How does this change the net charge on the soil colloid?

13. If there is a substitution of an [Al.sup.3+] for a [Si.sup.4+] in a tetrahedral sheet of a layer silicate clay, does the net negative charge of the clay increase or decrease?

14. If there is a substitution of an [Al.sup.3+] for a [Mg.sup.2+] in an octahedral sheet of a layer silicate clay, does the net negative charge of the clay increase or decrease?


Buseck, P. R. 1983. Electron microscopy of minerals. American Scientist 71: 175-185. (For a close-up examination of soil minerals.)

Dixon, J. B., and S. B. Weed. 1989. Minerals in the soil environment. Madison, WI: Soil Science Society of America. (A standard reference text in soil mineralogy.)


Allen, B. L., and B. F. Hajek. 1989. Mineral occurrence in soil environments. In J. B. Dixon and S. B. Weed (eds.), Minerals in soil environments, pp. 199-278. Madison, WI: Soil Science Society of America.

Bohn, H. L., B. L. McNeal, and G. A. O'Conner. 1985. Soil chemistry, 2nd ed. New York: John Wiley & Sons.

Brady, N. C., and R. R. Weil. 2002. The nature and properties of soil. Upper Saddle River, NJ: Prentice Hall.

Froedge, R. D. 1980. Soil survey of Christian County, Kentucky. Washington, DC: Dept. of Agriculture, Soil Conservation Service.

Gabler, R. E., R. J. Sager, and D. L. Wise. 1991. Essentials of physical geography, 4th ed. Philadelphia: Saunders College Publishing.

Plaster, E. J. 1997. Soil science & management. Clifton Park, NY: Thomson Delmar Learning .
TABLE 7-1 The most common elements in the Earth's crust.
(Gabler et al., 1991)

Element     %, by Weight   %, by Volume

O               46.6            93.8
Si              27.7             0.9
Al               8.1             0.5
Fe               5.0             0.4
Ca               3.6             1.0
Na               2.8             1.3
K                2.6             1.8
Mg               2.1             0.3

TABLE 7-3 Chemical composition of average igneous rocks, a slightly
weathered soil, and an intensively weathered soil. (Bohn et al., 1985)

                       Average        Slightly      Intensively
                      of Igneous   Weathered Soil    Weathered
                        Rocks       %, by Weight       Soil

Si[O.sub.2]               60             77             26
[Al.sub.2][O.sub.3]       16             13             49
[Fe.sub.2][O.sub.3]       7              4              20
Ti[O.sub.2]               1             0.6              3
MnO                      0.1            0.2             0.4
CaO                       5              2              0.3
MgO                       4              1              0.7
[K.sub.2]O                3              2              0.1
[Na.sub.2]O               4              1              0.3
[P.sub.2][O.sub.5]       0.3            0.2             0.4
S[O.sub.3]               0.1            0.1             0.3
TOTAL                   100.5          101.1           100.5

TABLE 74- Summary of mineralogic data of soils from various different
soil environments. (Derived from data in Allen and Hajek, 1989)

Soil Environment            Weathering     Dominant         Subordinate
                            Intensity      Mineral (a)      Minerals

Young sediments,
resistant parent materials   Mild          Smectite         Illite,
Hot and dry (arid)           Mild          Carbonates       Smectite,
Subhumid grassland           Intermediate  Smectite         Carbonates
Warm and humid forest        Intermediate  Smectite         Kaolinite,
Hot and humid forest         Strong        Kaolinite        Fe and Al
Tropical                     Strong        Fe and Al oxides Kaolinite

(a) In all cases the occurrence of mixed mineralogy is not uncommon.
These represent the generally recognized dominant minerals in many
such soils.

TABLE 7-5 Partial profile description of a soil from the Frondorf
series from Christian County, Kentucky, which formed in parent
materials of loess over sandstone residuum. (Adapted from Froedge,

                                       Primary    Secondary Clay
Horizon   Depth   Sand   Silt   Clay   Minerals      Minerals

           cm             %                     %

A          0-5    15.9   75.1    9.0      88            12
E          5-18   20.1   73.0    6.9      86            14
Bt1       18-33   14.5   60.1   25.4      63            37
Bt2       33-53   18.3   54.2   27.5      65            35
2Bt3      53-84   24.6   52.6   25.8      70            30
2R          >84
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Title Annotation:Section 3 Soil Chemical Properties
Publication:Fundamental Soil Science
Date:Jan 1, 2006
Previous Article:Chapter 6 Soil water, temperature, and gas relations.
Next Article:Chapter 8 Surface chemistry of soils.

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