The role of weathering processes in the formation of soil mineral composition

Weathering of rocks and the formation of soil minerals. Once rocks are exposed on the Earth's surface, they begin to break down under the influence of various factors. They undergo chemical changes, acquire new properties, and turn into loose products. The process of breaking down massive rocks and transforming them into loose products is called weathering. Weathering always, to one degree or another, precedes the soil formation process.

Due to the diversity of factors involved in the weathering of rocks, three forms are distinguished: physical, chemical, and biological. However, this division is arbitrary. In nature, all these forms of weathering are intertwined into a single and multifaceted process of pulverizing and transforming rocks and minerals.

The main factors of physical weathering are temperature, mechanical forces of water, wind, and glacier movement. They contribute to the transformation of rocks into smaller fragments, rubble, sand, and dust, crushing them mechanically without changing their petrographic and chemical composition. In such cases, it is easy to determine from the rock fragments what material this loose debris was formed from. Under the influence of physical weathering factors, massive crystalline rock becomes loose and permeable to water and air, but does not yet possess sufficient moisture retention. At the same time, as rocks are crushed, the surface area of contact between the fragments and the elements of the atmosphere and biosphere increases, which contributes to a more energetic impact of chemical and biological weathering factors upon them.

Chemical weathering is understood as the aggregate of phenomena resulting in the chemical decomposition of minerals that compose certain rocks, changing their chemical composition and structure, with new compounds—minerals—arising as a result of these processes.

The factors of chemical weathering are atmospheric water, carbon dioxide, and oxygen.

Water is an energetic solvent for rocks and minerals. The reactivity of water increases with rising temperature and its saturation with carbon dioxide. The latter gives water an acidic reaction, which increases its destructive effect on minerals. Under the influence of water with carbon dioxide and oxygen dissolved in it, a number of significant changes occur in the rock, resulting in the formation of a loose, fine mass that differs from the original rock.

The main reaction of water with rock minerals—hydrolysis—leads to the replacement of alkali and alkaline-earth metal cations in the crystal lattice with hydrogen ions from dissociated water molecules, i.e., exchange decomposition occurs. This reaction can be schematically considered using the example of orthoclase:

K2Al2Si6O16 + 2H2O → H2Al2Si6O16 + 2KOH

The base formed (KOH) causes an alkaline reaction of the solution, during which further destruction of the orthoclase crystal lattice occurs with the cleavage of part of the silica and the formation of kaolinite:

H2Al2 Si6O16 + nH2O →H2Al2 Si2O8H2O + 4SiO2.nH2O

In the presence of CO2, KOH transforms into the form of carbonates:

2KOH + CO2 →K2CO3 + H2O

Another process associated with water, which plays an important role in the chemical weathering of rocks, is hydration. This is a chemical process of attaching water particles to mineral particles to form compounds with properties different from the original ones. Thus, upon the hydration of hematite (Fe2O3), limonite (Fe2O3.3H2O) is formed:

2Fe2O3 +3H2O = 2Fe2O3. 3H2O

Hydration is also observed in silicates and aluminosilicates, which leads to their loosening and subsequently ensures closer interaction with water, air, and other weathering factors. Later, as a result of chemical reactions, kaolinite, halloysite, montmorillonite, nontronite, and sericite are formed from secondary aluminosilicates. These minerals have the form of small crystals and possess a number of fundamental properties inherent to colloids—exchange absorption capacity, high surface energy, etc. Due to these colloidal properties, these secondary minerals are of great agronomic importance.

Oxidation is also a widespread reaction in the weathering zone. Numerous minerals containing ferrous iron or other elements capable of oxidation are subject to this process. As an example of oxidation reactions during weathering, let us consider the interaction of sulfides with oxygen in an aqueous environment. It should be noted that the process of metal oxidation can only occur in the presence of moisture:

2FeS2 + 7O2 + 2H2O = 2FeSO4 + 2H2SO4

12FeSO4 + 6H2O + O2 = 4Fe2(SO4)3 + 4Fe(OH)3

2Fe2(SO4)3 + 9H2O = 2Fe2O3. 3H2O + 6H2SO4

During the oxidation of pyrite, along with sulfates and iron oxide hydrates, sulfuric acid is formed, which participates in the creation of new minerals.

The attachment of oxygen and water to the ferrous iron ions within the crystal lattice exerts a wedging effect on the lattice, contributing to its further destruction.

Thus, chemical weathering consists of the destruction of the mineral crystal lattice and the formation of compounds with different compositions and properties, the main causes of which are hydration, hydrolysis, and oxidation.

The mineral part of the soil is the basis of its soil fertility, which is formed during the weathering of rocks. For an agronomist, this process is important because it is during weathering that nutrients, originally locked in solid minerals, are released. Biological factors act simultaneously with physical and chemical ones, accelerating the destruction of the parent rock and creating conditions for plant development.

Microorganisms, plants, and livestock animals not only mechanically loosen the rock but also deeply change its chemical composition. The carbon dioxide released during respiration increases the acidity of the local environment and enhances the dissolving action of water. In addition, living organisms produce organic and mineral acids, which convert poorly soluble compounds into forms accessible to plants.

How living organisms turn barren rock into soil

Living organisms accelerate the destruction of rocks and convert elements into a form accessible to plants. The substances they release trigger various chemical reactions. Each type of microorganism and plant plays its own specific role in this process:

• Nitrifying and thionic bacteria release nitric and sulfuric acids, which dissolve strong mineral compounds.
• Silicate bacteria break down feldspars and phosphorites, releasing accessible potassium and phosphoric acid.
• Iron bacteria oxidize and break down iron compounds.
• Diatoms extract silicic acid from rocks, disrupting their structure.
• Mosses and lichens destroy rocks chemically (by releasing CO2 and acids) and mechanically (by penetrating the cleavage planes of minerals with hyphae and rhizoids).

Mineral stability and the influence of climate on soil composition

The rate of destruction depends directly on the mineralogical composition of the parent rock. Some minerals are destroyed quickly, while others remain unchanged for centuries. Weathering products are divided into primary, which retain their original composition (quartz, micas, feldspars), and secondary, which are synthesized anew under the influence of the environment.

• Most stable rocks — Metamorphic (e.g., quartzites)
• Least stable rocks — Sedimentary
• Most stable mineral — Quartz
• Average mineral stability — Feldspars
• Low mineral stability — Minerals with ferrous iron

The direction of soil formation and the composition of secondary minerals are determined by climatic conditions. The main factors here are air temperature and the volume of precipitation, which set the intensity of hydrolysis. Depending on the climate, weathering occurs in one of two types:

• Siallitic type (temperate climate, moderate precipitation) — predominantly secondary alumino- and ferrosilicates are formed.
• Allitic type (humid tropical climate) — intense hydrolysis occurs with the formation of silicon, aluminum, and iron oxide hydrates.

As a result of these processes, massive rock is transformed into loose weathered crust. It is a mixture of primary and secondary minerals enriched with nutrients. Such a change in the physical and chemical properties of the parent rock creates all the necessary conditions for the life of crop plants.

The number of primary minerals making up igneous rocks is in the hundreds; however, the number of minerals that make up the bulk of the rocks is small. Thus, according to F.W. Clarke, igneous rocks have the following mineralogical composition: feldspars – 59.5%, amphiboles (hornblende) and pyroxenes – 16.8%, quartz – 12.0%, micas – 3.8%, other minerals – 7.9%. These minerals, as noted earlier, are characterized by different resistance to weathering. In terms of strength, they can be arranged in the following sequence: quartz > feldspars > hornblende and pyroxenes > micas. In this regard, the relative content of primary minerals in soils and soil-forming rocks is different from that in igneous rocks. The predominant mineral in them is quartz, its share being 40–60%. Second place is occupied by feldspars (20% or more), which are characterized by high mechanical strength but are less resistant to chemical weathering. Amphiboles, pyroxenes, and micas also easily succumb to weathering, so their content in soils and loose rocks is small.

The resistance of primary minerals to weathering is determined by their chemical composition and crystal structure. The minerals under consideration have ionic-type structures formed by oppositely charged ions. The spatial arrangement of ions in mineral crystals is called a crystal lattice. It is characterized by a geometrically regular structure. Therefore, mineral crystals have the shape of regular polyhedra. It is individual for each mineral.

The mutual arrangement of anions and cations in a crystal lattice is determined by their volume or radius. The number of ions of opposite charge surrounding a given ion is called the coordination number. The latter determines the mutual arrangement of ions or the shape of the structural unit of the crystal lattice. Three main structural units are distinguished: the tetrahedron, the octahedron, and the cube. The tetrahedron is a three-dimensional figure with four triangular sides. Schematically, it can be represented as a pyramid with a triangular base. A silicon cation Si4+, less frequently aluminum Al3+, is located at the center of such a pyramid, while oxygen ions are at the vertices. The tetrahedron is a compact structure, with the distance between oxygen ions being 0.28 nm. The coordination number of cations located inside the tetrahedron is 4.

The octahedron can be viewed as two four-sided pyramids connected by their bases. It has eight sides and six vertices. Oxygen ions are located at the vertices. The space inside the octahedron is larger than in the tetrahedron, the bonds within it are longer and less stable. It contains medium-sized cations, such as aluminum Al3+, magnesium Mg2+, and iron Fe2+. The coordination number of octahedron cations is 6.

The cube is the simplest of the possible forms of oxygen ion arrangement. It provides space for cations such as Na+, Ca2+, and K+. The coordination number of such a structural unit is 8.

The main structural element of oxygen silicon compounds, which are widespread in the soil, is the silicate tetrahedron (SiO4)4–. It possesses four free valence bonds that can be compensated by the attachment of cations or by bonding with other silicate tetrahedra.

By connecting with each other via oxygen ions, tetrahedra form various types of structures: island, chain, ribbon, sheet (layered), and framework. The distribution of these structures in minerals is as follows:

• Framework: feldspars, quartz.
• Chain: pyroxenes.
• Sheet: micas, clay minerals.
• Ribbon: amphiboles.
• Island: olivine.

Island structures include silicate radicals containing 1, 2, 3, or more, but a finite number of tetrahedra (Fig. 55; Remezov N.P., 1957). In chain, ribbon, sheet, and framework structures, silicate tetrahedra form infinite radicals, linking into chains, ribbons, sheets, and frameworks (Fig. 56; Remezov N.P., 1957). The formula of such radicals shows the number of atoms in an elementary unit, which is repeated an infinite number of times in the structure.

Fig. 55. Island silicate radicals

Fig. 56. Silicate radicals: a – chain; b – ribbon; c – sheet

The significance of primary minerals is multifaceted: the agrophysical properties of the soil depend on their quantity; they serve as a reserve source of nutrients for plants, and secondary minerals are formed on their basis.

Secondary minerals in soils are represented by three main groups: simple salt minerals, hydroxide and oxide minerals, and clay minerals.

Simple salt minerals are formed during the weathering of primary minerals, as well as a result of the soil-forming process. They can accumulate in soils in sufficiently large quantities under arid climate conditions. Such minerals include: calcite CaCO3, magnesite MgCO3, dolomite [Ca, Mg] (CO3)2, soda Na2CO3·10H2O, gypsum CaSO4·2H2O, mirabilite Na2SO4·10H2O, halite NaCl, sylvine KCl, phosphates, nitrates, etc.

Hydroxide and oxide minerals are represented by the hydroxides of silicon, aluminum, iron, and manganese. They were formed during the weathering of primary minerals in an amorphous form as hydrated high-molecular-weight gels, then underwent a stage of dehydration and crystallization, forming oxides and hydroxides of a crystalline structure. Crystallization is facilitated by extreme temperatures (both high and low), drying, and oxidizing soil conditions. This group of minerals includes chalcedony and quartz, formed as a result of the transformation of silicon hydroxide SiO2·nH2O; pyrolusite MnO2 and psilomelane mMnO·MnO2·nH2O – derivatives of manganese hydroxide.

Secondary minerals: how clay determines soil properties

During the process of rock weathering and the crystallization of iron and aluminum hydroxides, secondary minerals are formed. These include boehmite (Al₂O₃·H₂O), hydrargillite or gibbsite (Al₂O₃·3H₂O or Al(OH)₃), hematite (Fe₂O₃), goethite (Fe₂O₃·H₂O), and hydrogoethite (Fe₂O₃·3H₂O). As the degree of crystallinity increases, the solubility of these compounds decreases.

Clay minerals — secondary alumino- and ferrosilicates — play a crucial role in the formation of the structure. They are formed through the chemical interaction of silica, sesquioxides, and bases (potassium, magnesium, and less frequently calcium and sodium). The most widespread groups in soils are montmorillonite (montmorillonite, nontronite, beidellite, saponite), kaolinite (kaolinite, halloysite, dickite), as well as hydromicas, vermiculite, and chlorites.

These secondary silicates determine the physical and chemical properties of soil. It is they that give the clay fraction its plasticity, viscosity, and swelling capacity. The key characteristics of clay minerals include:

• High dispersity (the size of plate-like or needle-shaped crystals does not exceed a few micrometers).
• Absorption capacity — the ability of cations on the surface of crystals to exchange for elements from the soil solution.
• Inconsistency of chemical composition due to the mutual substitution of elements in the crystal lattice (for example, silicon for aluminum, aluminum for iron or magnesium).
• Presence of chemically bound water, which is released only upon intense heating.

Particle-size distribution: the basis of the water-air regime

As a result of weathering, dense rocks turn into a loose mass consisting of particles of various sizes. Such a system is polydisperse, and its properties directly depend on the ratio of the fractions. Particle-size distribution determines the water, air, and thermal regimes of a field, as well as plant nutrient conditions. More than 80–90% of the mass of the solid soil phase consists specifically of mineral components.

All mechanical elements (isolated pieces of rocks, minerals, and amorphous compounds) are divided into two main categories. Particles larger than 1 mm make up the soil skeleton, while fractions smaller than this threshold are called fine earth. Within the fine earth, two key groups are distinguished, separated by a boundary of one-hundredth of a millimeter:

• Physical sand — fractions with a particle size of more than 0.01 mm.
• Physical clay — fine fractions with a particle size of less than 0.01 mm.

For a detailed assessment of soils and rocks by particle-size distribution, a classification of mechanical elements is used, grouping them by particle size.

Individual groups of mechanical elements influence soil properties in different ways. This is due to their differing mineralogical and chemical composition, and various physical and physicochemical properties. Therefore, numerous quantitative combinations of mechanical elements have led to a great diversity of soils based on particle-size distribution. These soils, in turn, are combined into several groups according to this indicator.

One of the first scientific classifications of soils by particle-size distribution was provided by N.M. Sibirtsev. It was binary and based on the ratio of physical clay to physical sand. Currently, the more advanced classification by N.A. Kachinsky is widely used.

According to this classification, the main division is made based on the content of physical clay and physical sand. All rocks and soils are grouped by particle-size distribution into 9 categories with physical, physicochemical, and chemical properties characteristic of each group. This classification is compiled taking into account soil types. This is because, as noted by N.A. Kachinsky (1958), "clay or another soil by mechanical composition is not a substrate containing in all cases x% physical clay and y% sand, but a soil possessing specific viscosity, resistance to ploughing, and other properties. And since soil properties depend not only on the quantity of mechanical elements of a certain size but also on the properties of these elements, and they are not identical in different soils, it is logical to differentiate classifications with respect to soil types."

In order to highlight the ratio of gravel, sand, silt, and clay fractions in the particle-size distribution, the concept of predominant fractions is introduced into the classification table. These are the gravelly fraction (1–3 mm), sandy (0.05–1 mm), coarse-silty (0.01–0.05 mm), silty (0.001–0.01 mm), and clayey (<0.001 mm).

The detailed classification takes into account the ratio of all the aforementioned mechanical element fractions. For example: chernozem soil contains 70% physical clay, 45% clay, 25% silt (medium and fine), 20% coarse silt, 10% sand. In this soil, the first predominant fraction is clay, the second in quantity is silt, and the third is coarse silt and sand. Such soil would be classified by particle-size distribution as: light clay, silty-clayey.

A brief scale, which takes into account only the ratio of physical clay to physical sand, is used for approximate determination of particle-size distribution in the field (using simplified tools and methods) and for compiling small-scale maps. A detailed scale is recommended for: mapping test plots and State Variety Trial plots; large-scale studies for soil reclamation purposes; surveying areas for special crops (beet, cotton, tea, vineyards, horticultural crops); and all soil surveys, starting from a scale of 1:25000 and more precise.

When using either of the two named scales, the classification of stony soils by particle-size distribution must be accompanied by indications of their stoniness degree. According to this indicator, the following soils are distinguished: 1) non-stony – less than 0.5% stony material, the soil is considered normal in terms of tillage; 2) slightly stony – 0.5–5.0% stony material; provided that this material consists of fine crushed stone or pebbles, the soil is tilled normally, but accelerated wear of the working surfaces of tillage implements is observed; 3) moderately stony – stony material content of 5–10%; for normal tillage, removal of large stony material is necessary; 4) strongly stony – >10% stony material; cultivation of annual crops requires labor-intensive work to remove stony material from the field.

Knowledge of the particle-size distribution of soils allows for predicting some of their properties to a certain extent. Thus, the more clayey a soil is, or, as is commonly said, the heavier it is in terms of particle-size distribution, the more dense, viscous, and less water-permeable it should be. Conversely, the more sand a soil contains, that is, the lighter it is in terms of particle-size distribution, the less viscous, less dense, and more water-permeable it will be. In addition, heavy soils are generally richer in humus, ash elements, and nitrogen, and are characterized by a higher cation exchange capacity.

However, this rule reflects the phenomenon only in a very general form. The presence of structure has a great influence on the listed soil properties. The same clay in a structureless state can be completely impermeable to water, while in the presence of a strong nut-like structure, it will transmit water well.

In addition to the particle-size distribution of soils, their aggregate composition is also distinguished. The fact is that in natural soils, individual mineral grains of which they consist, i.e., individual mechanical elements, are cemented into clumps of various sizes with a diameter ranging from a few millimeters to thousandths of a millimeter. These clumps are called aggregates. Aggregates are combined into fractions by their size in accordance with the same scale as for particle-size distribution. Based on this, the aggregate composition of soils is determined. The latter is of great importance for judging the properties of soils, as such aggregates are usually water-stable and remain in the soil without breaking down into their constituent grains.

By comparing the particle-size and micro-aggregate composition, one can find an objective numerical indicator that characterizes the soil's ability to form aggregates. This is the soil dispersion coefficient (Kd), proposed by N.A. Kachinsky. It represents the ratio of the content of silt particles in the soil, determined by micro-aggregate analysis (a), to their content determined by particle-size analysis (b), expressed as a percentage: a

Kd  100 %. b

The higher the dispersion coefficient, the more dispersed the soil is, and the worse its water-air properties and, as a consequence, the growing conditions for plants.