Weathering of Minerals

Last updated: July 24, 2026 Русский Español

Hello, dear listeners. Today we begin our discussion of the most amazing and mysterious process—how lifeless, solid rock transforms into the cradle of life, into soil. We will start at the very beginning, with the first step of this long journey, and answer the question: what forces set this mechanism in motion? How do millions of years of inanimate matter acquire a soul and become a fertile medium?

Our main goal is to understand the basic principles that will help us later navigate the infinite diversity of the world's soils. This entire long journey is called weathering. Weathering is not just destruction; it is a process of transformation that includes both the breakdown of old minerals and the synthesis of new ones, which determine the properties of the future soil (Birkeland, 1984; Weil & Brady, 2017). And before complex chemical reactions can begin, the rock must be prepared, crushed, so that every grain of sand and every tiny crystal is ready for change. This first preparatory role is played by physical weathering.

1. Physical Weathering: Breakers of the Hard Rock

Physical weathering is the process of mechanical disintegration of rock and minerals into smaller fragments without changing their chemical composition and mineralogical structure (Birkeland, 1984; Eash et al., 2016). Imagine taking a piece of rock and smashing it with a hammer—you changed its size, but not its composition. In nature, the role of such a hammer is played by various physical agents. Although the chemical composition does not change, the importance of physical weathering for soil formation is colossal: it increases the surface area of the rock, opening access for future chemical reactions and biological activity. In this sense, physical weathering is the great preparer (Foth, 1990).

Temperature Weathering (Thermoclasty)

One of the main agents of destruction is temperature fluctuations. During the day, the sun's rays heat the surface of the stone, and at night it cools. The minerals that make up the rock expand and contract at different rates and to different degrees (Birkeland, 1984; Eash et al., 2016). For example, dark minerals heat up and cool down faster than light ones. This creates internal stress at grain boundaries, and over time, repeated day after day for thousands of years, this stress leads to cracking and exfoliation of the rock.

This process is especially effective in regions with sharp daily temperature fluctuations, such as deserts, where the rock surface can heat up to 80 °C during the day and cool down to 20 °C at night (Scheffer et al., 2018). Such a process is sometimes called "exfoliation," when layers gradually peel off the rock surface, like the skin of an onion (Weil & Brady, 2017).

Frost Weathering (Cryoclasty)

An even more powerful destroyer is water that freezes. When freezing, water expands by about 9% and creates tremendous pressure (Birkeland, 1984). If water penetrates cracks, scratches, and pores of the stone, then upon freezing it acts like a wedge, widening these cracks. The process of repeated freezing and thawing causes huge boulders to split into many angular fragments. This process is most active in cold regions and high mountains, where freeze-thaw cycles are repeated dozens or hundreds of times a year (Weil & Brady, 2017). It is no coincidence that on many mountain peaks we see fields of rocky scree, which are a direct result of frost weathering.

Salt Weathering

In arid regions, where water often evaporates rather than seeps deep, salt weathering plays a huge role. Water contains dissolved salts. When it evaporates from the surface or from cracks in the rock, salts crystallize. Salt crystals grow and exert pressure on the walls of pores and cracks, which also leads to rock disintegration (Birkeland, 1984). This process, like frost action, is greatly amplified by repeated wetting and drying cycles (Scheffer et al., 2018). Salt weathering is one of the main factors creating bizarre weathering forms in deserts.

Other Factors

Physical weathering also includes abrasion—friction and grinding of particles against each other under the action of water, wind, or ice (Foth, 1990; Weil & Brady, 2017); destruction by tree roots, which, penetrating cracks, act as natural wedges (Birkeland, 1984); and even the effect of fire. During forest fires, the rock surface heats up rapidly and expands, while the inner layers remain cold, leading to exfoliation and cracks (Birkeland, 1984; Buol et al., 2011).

2. Chemical Weathering: The Rebirth of Matter

If physical weathering is the destroyer, then chemical weathering is the creator. It changes the very essence of minerals, their chemical composition and crystal lattice. It is here, in chemical reactions, that hard, inert rock begins to "come alive" and transform into that fine and active substance capable of retaining water, nourishing plants, and giving rise to life. The main agent of this process is water, but not just as a solvent, but as an active chemical reactant. The key reaction that triggers this rebirth is hydrolysis. We will examine it in more detail and see how feldspar is transformed into kaolinite.

2.1. Hydrolysis: The Heart of Chemical Weathering

Hydrolysis (from Greek hydor — water and lysis — decomposition) is a chemical exchange reaction between a mineral and water, in which hydrogen ions (H⁺) from water replace metal cations (K⁺, Na⁺, Ca²⁺, Mg²⁺) in the crystal lattice of the mineral (Eash et al., 2016; White, 2006). This process is the main "engine" that processes primary minerals into secondary ones.

Example: Transformation of Feldspar into Kaolinite

The classic and most striking example of hydrolysis is the transformation of orthoclase (potassium feldspar) into kaolinite. This is essentially the alpha and omega of weathering for many rocks. Let us break down this process step by step using the reaction equation (Foth, 1990; White, 2006; Weil & Brady, 2017):

Starting material: Orthoclase — KAlSi₃O₈

Reactants: Water (H₂O) and carbon dioxide (CO₂), which, dissolving in water, forms weak carbonic acid (H₂CO₃). Carbonic acid is the main supplier of H⁺ ions in the soil solution. It is formed as a result of respiration of plant roots and microorganisms (Buol et al., 2011).

Reaction products: The hydrolysis of orthoclase can be represented as a complex process, the overall scheme of which looks like this:

$$\underset{orthoclase}{2KAlSi_3O_8} + \underset{carbonic acid}{2H_2CO_3} + \underset{water}{9H_2O} \to \underset{kaolinite}{Al_2Si_2O_5(OH)_4} + \underset{silicic acid}{4H_4SiO_4} + \underset{potassium ions}{2K^+} + \underset{bicarbonate ions}{2HCO_3^-}$$

What happened in this equation?

1. Proton attack: Hydrogen ions (H⁺) from carbonic acid "attack" the crystal lattice of orthoclase. They displace potassium ions (K⁺), which pass into the soil solution. Potassium ions are a valuable plant nutrient that becomes available to roots (Foth, 1990).

2. Framework destruction: The loss of potassium makes the lattice unstable. Water continues to interact with the remaining aluminosilicate framework, breaking bonds between silicon and aluminum. This process is, in essence, the "unraveling" of the complex three-dimensional structure of feldspar.

3. Synthesis of a new mineral: The molecules of silicic acid (H₄SiO₄) and aluminum ions formed as a result of hydrolysis rearrange into a new structure. Silicon tetrahedra and aluminum octahedra join into layers, forming the crystal lattice of the clay mineral — kaolinite (Birkeland, 1984; Buol et al., 2011; Scheffer et al., 2018).

4. Formation of soluble products: Silicic acid and bicarbonates (HCO₃^-) are soluble compounds. They can either be removed from the soil by percolating water (leaching) or participate in other chemical reactions (Weil & Brady, 2017).

This process is the foundation of all soil formation. From insoluble, massive feldspar, finely dispersed, chemically active kaolinite is born. It is these clay minerals, along with sesquioxides, that will later become the main arena for all ion-exchange processes that determine soil fertility. Note that for each molecule of feldspar, many molecules of water and carbonic acid are required. Without water and biological activity (which supplies CO₂), this reaction would be impossible or would proceed thousands of times slower (Weil & Brady, 2017).

2.2. Other Types of Chemical Weathering

In addition to hydrolysis, other processes participate in chemical weathering:

Oxidation: This is the reaction of minerals with oxygen. The oxidation of iron (Fe²⁺ → Fe³⁺) is most important for soil formation. In primary minerals (e.g., olivine, biotite), iron is in the ferrous (divalent) form. In air, in the presence of water, it oxidizes to the ferric (trivalent) form and forms insoluble iron hydroxides (hematite — Fe₂O₃, goethite — FeOOH) (Scheffer et al., 2018; Weil & Brady, 2017). These compounds give most soils their characteristic yellow, orange, and red hues (Foth, 1990).

Hydration: The addition of water molecules to a mineral. For example, the transformation of anhydrous calcium sulfate (anhydrite) into gypsum (CaSO₄·2H₂O). Although this does not change the chemical formula radically, hydration increases the volume of the mineral, which can contribute to physical rock disintegration (White, 2006).

Dissolution: This is congruent dissolution, where the mineral passes entirely into solution without forming a new solid phase. A good example is the dissolution of calcite (CaCO₃) in water containing carbon dioxide:

$$CaCO_3 + H_2CO_3 \to Ca^{2+} + 2HCO_3^-$$

This process leads to the formation of karst caves and the destruction of limestone rocks (Birkeland, 1984; Eash et al., 2016). Gypsum (CaSO₄·2H₂O) and halite (NaCl) are also readily soluble.

Complexation (Chelation): This is a very important but often overlooked process. Organic acids released by microorganisms, lichens, and plant roots form stable water-soluble complexes—chelates—with metal ions (especially Fe and Al). These complexes "extract" metals from the crystal lattice of minerals, greatly accelerating their breakdown. This is a process where chemistry is closely intertwined with biology—a bridge to our next section (Birkeland, 1984; Weil & Brady, 2017).

We have seen how chemistry, with the help of water and acids, reshapes the mineral world, transforming massive crystals into the finest clays. But the engine of this process, the source of that very carbonic acid and complex organic acids, is life. Living organisms not only use the products of weathering but also actively direct it.

3. Biological Weathering: Life as an Architect

If physical weathering is the hammer, and chemical weathering is water and acid, then biological weathering is the skilled master who combines their efforts and directs them in the right direction. Biological weathering is the transformation of minerals and rocks under the influence of living organisms and their metabolic products (Mukha et al., 2003). It is life that makes the process of soil formation irreversible and directed.

Biological weathering is not a separate process, but a powerful catalyst that accelerates both physical and chemical weathering by orders of magnitude. Without life on Earth, soil as we know it would simply not exist. As the outstanding scientist V. I. Vernadsky noted, weathering processes, especially in the biosphere, are always biogenic and bio-inert, and microscopic life plays a leading role in them (Mukha et al., 2003).

3.1. Plants: Roots as Wedges and Chemical Factories

The role of plants in weathering is multifaceted.

Physical impact: Roots of trees and large grasses penetrate microcracks in the rock. As the root grows, it thickens, and the pressure it exerts on the crack walls can reach 1–1.5 MPa (Scheffer et al., 2018). This is comparable to the pressure of freezing water! Roots act as natural wedges, pushing apart and splitting even the hardest rocks.

Chemical impact: Plant roots release into the rhizosphere—the zone around the root—a wide range of organic acids (citric, malic, oxalic) and complex organic chelators (Birkeland, 1984). These substances actively dissolve minerals, extracting necessary nutrients. In addition, roots, like all living organisms, respire, releasing carbon dioxide (CO₂), which, dissolving in water, forms carbonic acid—the main agent of hydrolysis (Buol et al., 2011). The concentration of CO₂ in soil air can be tens and hundreds of times higher than in the atmosphere, so soil solutions are much more aggressive toward minerals.

3.2. Microorganisms: Invisible Armies

Bacteria, fungi, actinomycetes, and algae are the invisible but most numerous and active participants in weathering.

Acid production: Microorganisms are powerful producers of organic and inorganic acids. For example, some bacteria oxidize sulfur to sulfuric acid and nitrogen to nitric acid, creating an extremely aggressive environment for minerals (Eash et al., 2016; Mukha et al., 2003). Fungi secrete citric and oxalic acids, which actively dissolve silicates.

Complexation: Microorganisms excrete complex organic substances that form stable water-soluble complexes—chelates—with metal ions (Fe, Al, Mn). These chelates "extract" metals from the crystal lattice, destroying the mineral from within (Birkeland, 1984; Weil & Brady, 2017).

3.3. Lichens: Pioneers of Life

Lichens are symbiotic organisms consisting of a fungus and an alga (or cyanobacterium). They are the first to colonize bare rocks and become true "pioneers" of weathering (Eash et al., 2016).

Physical breakdown: Fungal hyphae mechanically penetrate the rock, loosening its surface (Mukha et al., 2003).
Chemical attack: Lichens produce specific organic acids (e.g., oxalic acid) that actively dissolve minerals. Studies show that weathering crusts under lichens are significantly thicker and chemically more altered than on uncolonized surfaces (Birkeland, 1984). For example, under a lichen layer, there may be a sixfold enrichment in iron (see comparison of lichen and non-lichen weathering in Birkeland, 1984, Table 3-3).

3.4. Soil Fauna: Mixing and Acceleration

Animals living in the soil—earthworms, ants, termites, rodents—act as "ameliorators" (White, 2006; Weil & Brady, 2017).

Bioturbation: They actively mix the soil mass, transporting particles from lower horizons to upper ones and vice versa. This mixing promotes contact of fresh mineral particles with organic acids, water, and oxygen, greatly accelerating all weathering processes.
Aeration: Burrowing activity creates channels and pores, improving soil aeration and facilitating the penetration of water and air deep into the soil (Scheffer et al., 2018).
Comminution: By passing soil and plant residues through the digestive tract, worms and other invertebrates further grind mineral particles, increasing their surface area for chemical reactions.

We have seen how life not only uses the results of weathering but actively directs and accelerates it, combining physics and chemistry into a single process. With the appearance of life on Earth, weathering became a biogeochemical process.

Now that we know how minerals are broken down and transformed, a logical question arises: how fast does this happen? And why do some minerals disappear in thousands of years, while others remain virtually unchanged for millions of years? The answers to these questions will be found in the following sections of our lecture, devoted to weathering rates and Goldich's stability series.

4. Process Rates: Time as the Great Creator

Time is the fifth and perhaps the most mysterious factor of soil formation. We cannot see it, but we can measure its work by how far the transformation processes have gone. The rate of weathering is not a constant. It is a quantity that depends on many variables, and understanding these dependencies is the key to understanding the soil cover of the Earth (Jenny, 1941; Weil & Brady, 2017).

4.1. What Does the Weathering Rate Depend On?

Imagine a chemical reaction. Its rate depends on temperature, concentrations of reactants, surface area, and the presence of catalysts. Exactly the same applies to weathering:

Climate (Temperature and Humidity): These are the main regulators. In chemistry, there is an empirical rule: with a temperature increase of 10 °C, the rate of most chemical reactions increases by 2-3 times (Weil & Brady, 2017). That is why in hot and humid tropics, weathering proceeds orders of magnitude faster than in the cold Arctic or arid desert. Water is the universal solvent and reactant. Without water, hydrolysis and oxidation reactions practically stop.

Rock properties (Mineral composition and structure): This is our "Goldich stability series," which we will discuss in more detail in the next chapter. Some minerals, like quartz, are practically unaffected chemically, while others, like olivine, "burn out" in just a few thousand years (Birkeland, 1984; White, 2006). Moreover, porous, fractured rock has a huge surface area for reactions compared to massive, dense rock.

Surface area: Physical weathering prepares the material, increasing its surface area. Imagine a cube with a side of 1 cm. Its area is 6 cm². If we crush this cube into 1 mm cubes, the total surface area will increase to 600 cm²! By the same factor, the rate of chemical reactions will increase (Foth, 1990). Therefore, physical and chemical weathering processes always act in close conjunction, enhancing each other.

Presence of living organisms: As we have already discussed, microorganisms, plants, and their exudates are powerful catalysts. They supply acids, chelators, and increase CO₂ concentration hundreds of times, sharply accelerating all reactions.

Time: Weathering rate is not constant over time. It is usually highest in the initial stages, when fresh, reactive material is just beginning to be exposed. Over time, as secondary minerals and protective coatings form on the surface, the rate may slow down, approaching the so-called steady state—a state of dynamic equilibrium when the rate of formation of new minerals equals the rate of their destruction (White, 2006; Weil & Brady, 2017).

4.2. Quantitative Estimates of Rates

How can we measure this rate in real numbers? There are several approaches.

River water chemistry: This is one of the most popular methods. By analyzing the composition of water draining from a catchment basin and knowing the basin area, one can calculate how many tons of dissolved substances are removed from the territory per year. This provides an integrated estimate of the rate of chemical denudation (Birkeland, 1984). For example, studies show that in mountains where processes are intense, chemical denudation can reach tens of tons per square kilometer per year (Strakhov, 1967).

Data on rates for specific minerals: Attempts have been made to estimate how many years it would take for a 1 mm crystal to completely dissolve for different minerals (Birkeland, 1984, Table 3-12). Here are these figures (Lasaga, 1984):

Olivine: ~100 years
Diopside (pyroxene): ~6,800 years
Albite (Na-feldspar): ~80,000 years
Orthoclase (K-feldspar): ~520,000 years
Muscovite (mica): ~2,700,000 years
Quartz: ~34,000,000 years

These figures clearly show why olivines and pyroxenes disappear first in soils, while quartz remains even in the most ancient and weathered strata. Note the difference in rate between feldspars and micas.

Field chronosequences: This is a classic method in soil science. Researchers find a series of soils formed on the same material and under the same climatic conditions but of different ages (e.g., river terraces). By studying the degree of development of soil horizons (e.g., clay accumulation in the Bt horizon) on these terraces, one can construct a curve of weathering degree versus time (Weil & Brady, 2017). Studies in California (Arkley, 1964, cited in Foth, 1990) showed that a noticeable increase in clay content in the podzolic horizon requires tens of thousands of years. For example, a 1000-year-old soil had no signs of clay illuviation, while a 10,000-year-old soil already had 1.4 times more clay in the B horizon than in the A. This shows that clay formation is a very slow geological process.

Kinetic modeling: Modern scientists use complex kinetic equations to describe the dissolution rate of minerals. These equations take into account many factors: specific surface area, temperature, pH of the solution, presence of inhibitors, etc. (Scheffer et al., 2018). One such equation shows that the dissolution rate of a mineral is directly proportional to its degree of undersaturation in the solution. That is, the faster we remove weathering products (e.g., by flowing water), the faster the reaction will proceed.

4.3. General Conclusion

Weathering rates are extremely variable. They can range from several millimeters per year for readily soluble minerals in the tropics to a millimeter per thousand years for quartz in the dry Arctic.

Key ideas of this section:

1. Climate is the main conductor: Heat and moisture greatly accelerate all processes.

2. Mineral determines fate: Easily soluble minerals disappear first, leaving behind resistant quartz and secondary clays.

3. Time is the decisive factor: Even slow processes over millions of years create colossal changes, forming deep weathering crusts.

4. Life is an accelerator: The biological factor can increase weathering rates by orders of magnitude.

We now know that time plays a huge role, but it acts selectively. Some minerals disappear almost instantaneously (in geological terms), others remain almost eternal. This selectivity is the key to understanding the mineral composition of any soil.

Which mineral will disappear first, and which will remain in the end? The answer to this question is given by Goldich's stability series, which we will discuss in the next chapter.

5. Goldich's Stability Series: Who Survives in the Soil World?

Imagine that rock is a complex community of minerals. When it comes to the Earth's surface conditions, a struggle for survival begins. Some minerals "surrender" almost immediately, others resist for millions of years. This hierarchy of stability was described by American petrologist Samuel Goldich in 1938. His stability series is essentially the inverse of Bowen's reaction series, which describes the order of crystallization of minerals from magma (Goldich, 1938; Buol et al., 2011; White, 2006). That which crystallizes last at high temperatures turns out to be the most stable on the Earth's surface.

5.1. Structure of Goldich's Series

Goldich's series has the shape of an inverted Y (or, if viewed from above, a Y) and is divided into two main branches (Birkeland, 1984; Buol et al., 2011):

Left branch (Mafic—dark-colored):

Here are minerals rich in iron and magnesium, which crystallize at the highest temperatures.

Least stable — Olivine: (Mg,Fe)₂SiO₄. This is an inosilicate where silicon tetrahedra are not directly linked to each other but are held only by weak bonds with magnesium and iron ions. Therefore, olivine is destroyed first (Birkeland, 1984; Foth, 1990).

Pyroxenes (e.g., Augite): These are chain silicates. Silicon tetrahedra are connected into chains, but between the chains there are weak bonds with cations. More stable than olivine, but still very easily weathered (White, 2006).

Amphiboles (e.g., Hornblende): These are double-chain silicates, where two chains are joined together. Bonds within the ribbons are stronger, so stability increases (Foth, 1990).

Biotite: This is a mica, a layered silicate. Although it is already a sheet silicate, biotite, as we will see, has its own peculiarity. In the classic Goldich series, it stands above (less stable) compared to the next branch. However, as we will discuss below, modern data adjust its position (Buol et al., 2011).

Right branch (Feldspathic—light-colored):

Here are feldspars, which crystallize at lower temperatures.

Anorthite (CaAl₂Si₂O₈): Calcium feldspar. Contains a lot of calcium, which is easily leached. Feldspars have a framework structure, but the presence of aluminum and calcium ions makes them vulnerable (Birkeland, 1984).

Albite (NaAlSi₃O₈): Sodium feldspar. More stable than anorthite, since sodium is less susceptible to hydrolysis than calcium, and the aluminum content is lower (White, 2006).

Orthoclase (KAlSi₃O₈): Potassium feldspar (microcline). The most stable of the feldspars. Potassium is firmly held in the structure, and aluminum tetrahedra make up only a quarter of all (Birkeland, 1984; Foth, 1990).

Junction of branches:

Muscovite: Light mica. This is a very stable sheet silicate containing no iron. In Goldich's series, it stands at the boundary of the two branches, as a link, and is one of the most stable minerals after quartz.

Crown of the series:

Quartz (SiO₂): The most stable. This is a framework silicate where every oxygen atom is bonded to two silicon atoms. This structure is maximally strong, and chemical attack on quartz is minimal. That is why quartz dominates in sands and is the main mineral in the most weathered soils (Foth, 1990; Weil & Brady, 2017).

5.2. Why This Sequence?

Stability is determined by two key factors (Foth, 1990; White, 2006):

1. Degree of connectivity of silicon-oxygen tetrahedra: The more SiO₄ tetrahedra are connected to each other through shared oxygen atoms, the stronger and more stable the structure. In olivine they are not connected at all, in quartz they are fully connected. Accordingly, structural strength increases from isolated tetrahedra to the three-dimensional framework.

2. Strength of bonds between cations and oxygen: The strongest bond is Si–O, then Al–O, while bonds with alkali and alkaline-earth metals (K, Na, Ca, Mg) are significantly weaker. Therefore, minerals rich in calcium and magnesium are destroyed more easily than those where silicon predominates.

5.3. Important Clarification: Biotite and Vermiculite

In the classic Goldich series, biotite stands at the boundary with amphiboles. However, modern studies using X-ray diffraction have shown that biotite can very quickly transform into vermiculite. This reaction is essentially a transformation, not complete destruction. The 2:1 structure is preserved, but the composition of the interlayer space changes. Therefore, biotite should be considered less stable than previously thought, and in the series it is often placed before hornblende (Buol et al., 2011; Birkeland, 1984). That is, biotite may "disappear" as a mineral, but give rise to another layered silicate.

5.4. Practical Significance of Goldich's Series

This series is not just an academic classification. It allows soil scientists to literally read the history of weathering:

Young soils will contain the full spectrum of minerals, including olivine, feldspars, and biotite.
Soils of intermediate age lose olivine, pyroxenes, and amphiboles. Feldspars, biotite, and muscovite remain in them.
Ancient soils are depleted in almost all primary minerals except quartz, and secondary minerals begin to dominate, which we will discuss in the next chapter.

Now we know which mineral will "survive" the weathering process and which will be destroyed to give birth to something new. And this new stage—the formation of secondary minerals—is the culmination of the entire process. From fragments and ions, clays are born—those very particles that determine soil fertility.

What are these secondary minerals? How are they structured and why are they so important? We will discuss this in the next, sixth chapter of our lecture.

6. Secondary Minerals: New Birth from the Old

Secondary minerals are minerals that form on the Earth's surface as a result of chemical weathering of primary minerals, either by their transformation or by neoformation (synthesis) from their breakdown products (Foth, 1990; Buol et al., 2011). This is essentially a new mineral world that arises from the wreckage of the old. Secondary minerals are much finer than primary ones (usually < 0.002 mm, i.e., belonging to the clay fraction), and it is they that give soils their unique properties: plasticity, swelling, high sorption capacity (Scheffer et al., 2018; Weil & Brady, 2017).

6.1. Two Pathways for the Formation of Secondary Minerals

There are two main mechanisms for the birth of secondary minerals (Birkeland, 1984; Buol et al., 2011):

1. Transformation: This is the "soft" pathway. The original primary mineral retains its basic crystal structure (usually layered), but significant changes occur within it: removal or replacement of cations in the interlayer space, oxidation of iron, etc. A classic example is the transformation of biotite into vermiculite. During weathering, potassium is removed from the interlayer spaces of biotite, and iron is oxidized, leading to lattice expansion and the formation of vermiculite (Birkeland, 1984; Buol et al., 2011). This process does not require complete dissolution of the mineral; it occurs "in situ."

2. Neoformation (synthesis): This is the "radical" pathway. The primary mineral is completely dissolved, and its ionic components (Si, Al, Fe, Mg, Ca, etc.) enter the soil solution. From this solution, under certain conditions (pH, concentration, temperature), new, completely different minerals begin to crystallize. This is exactly how kaolinite is born from the hydrolysis products of feldspar. Neoformation is the main process for most clay minerals in soils (White, 2006).

6.2. Main Groups of Secondary Minerals

Secondary minerals can be divided into several main groups, each with its own structure and significance.

Clay Minerals (Silicate Clays)

This is the most important group, and they are the product of hydrolysis of silicates. All of them belong to phyllosilicates—layered silicates. Their structure is based on two types of layers (Buol et al., 2011; Scheffer et al., 2018):

Tetrahedral layer: Consists of silicon-oxygen tetrahedra (SiO₄) connected into an infinite flat grid.
Octahedral layer: Consists of octahedra, in the center of which is an aluminum (Al), magnesium (Mg), or iron (Fe) atom, surrounded by hydroxyl groups (OH).

The combination of these layers gives different types of clay minerals:

Kaolinite group minerals (1:1): Kaolinite and halloysite. Consist of one tetrahedral and one octahedral layer. Layers are strongly bonded by hydrogen bonds, so kaolinite does not swell, has a small specific surface area, and low cation exchange capacity (5-15 cmol(c)/kg). Kaolinite is the end product of weathering of many silicates in acidic and well-drained soils (Birkeland, 1984; Buol et al., 2011; Foth, 1990).

Hydromica group (2:1): Illite (hydromuscovite). Structure 2:1—two tetrahedral layers with an octahedral one between them. In the interlayer space are potassium ions (K⁺), which strongly bind the layers, so illite does not swell or swells weakly, but has a high cation exchange capacity (20-40 cmol(c)/kg). Illite is a product of incomplete weathering of micas and is widespread in young soils and soils of temperate climates (Scheffer et al., 2018; White, 2006).

Vermiculite group (2:1): Vermiculite. Structure 2:1, but in the interlayer space are hydrated cations (Ca²⁺, Mg²⁺), which are held more weakly than K⁺. Therefore, vermiculite swells, but moderately, and has a very high cation exchange capacity (100-150 cmol(c)/kg). It is a product of weathering of biotite and chlorite (Buol et al., 2011; Foth, 1990).

Smectite group (2:1): Montmorillonite, beidellite. Structure 2:1, but the interlayer space is filled with hydrated cations and a large amount of water. The bond between layers is very weak, so smectites swell strongly, have a huge specific surface area, and very high cation exchange capacity (80-120 cmol(c)/kg). Smectites are characteristic of soils with alkaline reaction, poor drainage, and rich in bases (Birkeland, 1984; Buol et al., 2011; Weil & Brady, 2017).

Chlorite group (2:1:1): Chlorite. Structure 2:1, but in the interlayer space there is a hydroxide layer, similar to brucite Mg(OH)₂ or aluminum hydroxide. This makes chlorite non-swelling and with moderate cation exchange capacity. It occurs both as primary (in metamorphic rocks) and as secondary (weathering product). In soils, so-called soil chlorites (or hydroxy-interlayered minerals) are often formed, when aluminum hydroxides enter the interlayer space of smectites and vermiculites (Birkeland, 1984; Scheffer et al., 2018).

Amorphous Minerals

This is a group of minerals that do not have a strict crystal lattice. They have a huge specific surface area and high reactivity.

Allophane and imogolite: These are aluminosilicates with short-range order. Allophane appears as hollow spherical particles (diameter 3-5 nm), and imogolite as tubes. They form during rapid weathering of volcanic ash, as well as in podzolic soils. Their cation exchange capacity is strongly pH-dependent (variable charge) (Buol et al., 2011; Scheffer et al., 2018).

Oxides and Hydroxides

This is a very important group that forms as a result of oxidation and hydrolysis of iron, aluminum, manganese, and silicon.

Aluminum: Gibbsite (Al(OH)₃). This is the main mineral in strongly weathered tropical soils (ferrallitic soils). It forms when the concentration of silicon in the solution is very low, and aluminum precipitates as hydroxide (Birkeland, 1984; Buol et al., 2011).

Iron: Goethite (α-FeOOH) and hematite (α-Fe₂O₃). Goethite (yellow-brown) and hematite (red) are the most common iron minerals in soils. They determine soil color and participate in particle aggregation. Ferrihydrite (Fe₅HO₈·4H₂O) is a poorly crystallized, amorphous iron hydroxide that is often the first product of iron oxidation and is found in young soils and podzols (Scheffer et al., 2018; Weil & Brady, 2017).

Manganese: Birnessite, vernadite, etc. Manganese oxides often occur as black spots and concretions. They have a high capacity for sorption of heavy metals.

Silicon: Opal (SiO₂·nH₂O). This is amorphous silica, which can form in soils of arid regions, as well as in the form of phytoliths (plant opal bodies) (Birkeland, 1984).

6.3. Patterns of Secondary Mineral Formation (Stability and Conditions)

The formation of a particular secondary mineral is governed by strict thermodynamic laws. Two factors play a key role: the concentration of silica (H₄SiO₄) in the soil solution and pH (Birkeland, 1984; Buol et al., 2011; White, 2006).

High silicon (Si) concentration and alkaline reaction: 2:1 minerals—smectites (montmorillonite)—are formed. This occurs under poor drainage, in arid regions, where removal of weathering products is hindered.

Medium silicon concentration and acidic reaction: 1:1 minerals—kaolinite—are formed. This is the classic situation for humid (wet) regions with good drainage, where silicon is actively leached but not completely.

Low silicon concentration (less than 10⁻⁵ mol/L) and acidic reaction: Gibbsite (Al(OH)₃) is formed. This is the desilication stage, characteristic of ancient, intensely weathered tropical soils.

This pattern is clearly demonstrated on mineral stability diagrams (see, e.g., Buol et al., 2011, Figure 4.10; Birkeland, 1984, Figure 2.26). Understanding these relationships allows predicting the mineralogical composition of soil, and hence its properties.

6.4. Significance of Secondary Minerals for Soil Properties

It is secondary minerals that determine the most important agronomic properties of soil:

1. Sorption capacity (Cation Exchange Capacity — CEC): Clay minerals and humus are the main carriers of exchangeable cations (Ca²⁺, Mg²⁺, K⁺, Na⁺, H⁺, Al³⁺). The higher the CEC, the more nutrients can be held in the soil and be available to plants (Foth, 1990; Weil & Brady, 2017). 2:1 minerals (smectite, vermiculite) have much higher CEC than kaolinite.

2. Physical properties (Structure, water regime): Swelling clays (smectites) greatly increase in volume when wet, and when dry produce deep cracks. This affects soil density, air and water regimes. Kaolinite, on the other hand, does not swell and forms more loose, permeable soils (Scheffer et al., 2018; Weil & Brady, 2017).

3. Fertility: Soils with high content of 2:1 minerals (e.g., Chernozems) are often more fertile because they better retain potassium and other cations. Strongly weathered soils with kaolinite and gibbsite are generally poor in bases and require fertilization (Birkeland, 1984; Foth, 1990).

Thus, secondary minerals are not just destruction products. They are new building blocks from which soil fertility is created. They form according to strict laws, depending on climate, parent material, and time, and it is their composition that determines what the soil will be like.

Now that we know what the second, "living" half of the solid phase of soil consists of, we can summarize. Our entire lecture was devoted to one process—the transformation of rock into soil. And the main visible result of this process is clay formation. In the concluding chapter, we will synthesize and show how all stages of weathering—from physical fragmentation to the synthesis of secondary minerals—merge into a single, continuous process of clay formation, which is the quintessence of the transformation of solid rock into soil.

7. Clay Formation: How Fertility Is Born

We have reached the most important and beautiful moment of our lecture. We have seen how physics crushes stone, how chemistry rearranges its structure, how life directs and accelerates these processes. Now we will see how all this unites in a single, continuous, and grand process—clay formation. It is the formation of clays that is the visible result that turns lifeless rock into soil—a complex, active, fertile medium (Birkeland, 1984; Foth, 1990; Weil & Brady, 2017).

7.1. What Is Soil Clay and Why Is It So Important?

Soil clay is not just fine particles. It is a special category of mineral particles less than 0.002 mm (2 microns) in size, which have fundamentally different properties from coarser sand and silt (Buol et al., 2011; Scheffer et al., 2018). Clay is the "workhorse" of the soil. It is clay that:

Retains water: The huge specific surface area and capillary forces allow clay to hold a significant amount of moisture available to plants (Foth, 1990; Weil & Brady, 2017).
Binds nutrients: Clay minerals have a negative charge and are able to retain cations (K⁺, Ca²⁺, Mg²⁺, NH₄⁺) on their surface, which can then be used by plant roots. This phenomenon is called cation exchange (Birkeland, 1984; White, 2006).
Forms soil structure: Clay particles, combining with organic matter and iron oxides, form aggregates—crumbs—that give the soil porosity, providing air and water access to roots (Scheffer et al., 2018).

Thus, clay formation is not just a change in mineral composition; it is the creation of the foundation for the entire soil ecosystem.

7.2. Mechanisms of Clay Formation

Clay minerals in soil can be formed by two main pathways: transformation and neoformation (synthesis). We have already briefly mentioned both processes, but now we will consider them in more detail.

Transformation: "Soft" Reorganization

Transformation is a process in which the primary mineral changes without complete destruction of its crystal lattice (Birkeland, 1984; Buol et al., 2011). It is like renovating a house: you change the interior decoration and utilities, but the walls and roof remain the same.

Classic example: Biotite → Vermiculite

Biotite is a dark mica (trioctahedral 2:1 mineral) containing potassium ions (K⁺) in the interlayer space. During weathering, the following changes occur:

1. Oxidation of iron: Iron in the octahedral layer is oxidized (Fe²⁺ → Fe³⁺). This increases the positive charge of the layer, weakening the bond with interlayer potassium (Birkeland, 1984; Scheffer et al., 2018).

2. Potassium displacement: Weakened potassium is replaced by hydrated cations (Ca²⁺, Mg²⁺) from the soil solution (Buol et al., 2011).

3. Lattice expansion: Hydrated cations are larger than K⁺, so the interlayer distance increases from 1.0 nm (biotite) to 1.4 nm (vermiculite) (Foth, 1990; Scheffer et al., 2018).

As a result, the 2:1 structure is preserved, but the mineral acquires new properties: it becomes swelling, its cation exchange capacity increases sharply. This process demonstrates how even small chemical changes can radically transform mineral properties.

Other examples of transformation:

Muscovite → Illite: A similar process, but slower, as muscovite (dioctahedral mica) is more stable.
Chlorite → Vermiculite/Smectite: Removal of the interlayer hydroxide interlayer from chlorite.
Smectite → Soil chlorite: Introduction of aluminum hydroxides into the interlayer space of smectite in an acidic environment (Buol et al., 2011).

Neoformation (Synthesis): "Rebirth"

This pathway is more radical. The original primary mineral is completely destroyed (dissolved) by hydrolysis, and from its breakdown products—ions and molecules—a completely new mineral is synthesized (White, 2006). This is no longer renovation, but building a new building from the bricks of the old one.

Key example: Feldspar → Kaolinite

We have already examined this reaction in detail in the chapter on chemical weathering. The hydrolysis of feldspar leads to all its components (K⁺, Si⁴⁺, Al³⁺) entering solution. When the concentration of silicon and aluminum in the solution reaches a certain level, they begin to crystallize into a new mineral structure—kaolinite (1:1 layer, dioctahedral) (Birkeland, 1984; Foth, 1990).

Conditions for the synthesis of various clay minerals:

Which particular clay mineral is synthesized depends on the chemical composition of the solution, pH, and drainage conditions (Buol et al., 2011; Weil & Brady, 2017):

Smectite (2:1): Forms in alkaline environments, at high concentrations of silicon, magnesium, and bases (Ca, Na). This is characteristic of poorly drained soils or in arid regions.
Kaolinite (1:1): Forms in acidic environments, at medium silicon concentration. This is the "golden mean"—well-drained soils with moderate silicon removal.
Gibbsite (Al(OH)₃): Forms in very acidic environments, when the silicon concentration is extremely low (less than 10⁻⁵ mol/L). This is the stage of complete desilication, characteristic of ancient, strongly weathered tropical soils (ferrallitic soils) (Birkeland, 1984; Buol et al., 2011).

7.3. Formation of Oxides and Hydroxides: Iron and Aluminum

In addition to clay silicates, oxides and hydroxides of iron, aluminum, and manganese are widespread in soils. Their formation is closely related to redox processes (Scheffer et al., 2018; Weil & Brady, 2017).

Iron:

In primary minerals, iron is in the divalent form (Fe²⁺). During weathering, it oxidizes to the trivalent form (Fe³⁺).

Then Fe³⁺ hydrolyzes to form sparingly soluble hydroxides and oxides:

Ferrihydrite (Fe₅HO₈·4H₂O): Amorphous, poorly crystallized. Often the first phase formed during rapid oxidation. Found in young soils and podzols (Birkeland, 1984).
Goethite (α-FeOOH): Yellow-brown, crystalline. The most common iron mineral in soils. Forms during slow oxidation (Scheffer et al., 2018).
Hematite (α-Fe₂O₃): Red, crystalline. Forms in warm and dry conditions, often from ferrihydrite. It is hematite that gives tropical red soils their characteristic color (Buol et al., 2011).

Aluminum:

When silicon is strongly leached from aluminosilicates (desilication), gibbsite (Al(OH)₃) forms—the main aluminum mineral in strongly weathered soils (Birkeland, 1984; Buol et al., 2011).

7.4. A Unified Process: How It All Works in Nature

Now we can draw a complete picture. Imagine a granite rock consisting of quartz, feldspar, and biotite:

1. Physical weathering (frost, temperature) creates cracks, breaking the rock.

2. Biological weathering (lichens, roots, microorganisms) penetrates the cracks, releasing acids and chelators.

3. Chemical weathering (hydrolysis, oxidation) begins to destroy primary minerals. Potassium is leached from feldspar, the structure loosens. Potassium is leached from biotite, iron oxidizes.

4. Transformation: Biotite, retaining its layered structure, turns into vermiculite. This process proceeds faster than the complete destruction of feldspar (Buol et al., 2011).

5. Neoformation: In an acidic environment, where silicon is actively leached, kaolinite is synthesized from the hydrolysis products of feldspar. This process is slower than biotite transformation, but as weathering deepens, it becomes dominant (White, 2006).

6. Oxidation: Iron from biotite and other minerals oxidizes, forming goethite and hematite, which color the soil in brown and red tones.

7. Result: As the process continues, first the least stable minerals disappear (olivine, pyroxenes), then feldspars and biotite, and eventually only the most stable primary minerals (quartz) and secondary clay minerals (kaolinite, vermiculite, smectite) and oxides (goethite, hematite, gibbsite) remain in the soil (Birkeland, 1984; Foth, 1990).

Concluding Remarks: Such Is the Story of the Transformation of Stone into Soil

Here is the complete picture. From the same parent rock, depending on climate, relief, and time, completely different soils can form.

In cold and dry climates, the process will stop at the stage of physical fragmentation—we will get coarse, skeletal soils.
In a temperate humid climate, where there is both heat and water, the full cycle is activated—and we get fertile soils with a high content of illite, vermiculite, and kaolinite.
In a hot and humid tropical climate, with enormous amounts of time, the processes go so far that only quartz, kaolinite, and iron and aluminum oxides remain from the original rock—red, deeply weathered, but often nutrient-poor soils.

This path from stone to clay is the fundamental process that underlies all soil formation. By understanding it, we can not only describe but also predict soil properties, and therefore manage them wisely.

References

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