Elementary Soil-Formation Processes

Introduction: From Factors to Processes

In previous lectures, we have already become acquainted with the fundamental concept of soil-forming factors—climate, relief, parent material, organisms, and time (Jenny, 1941; Buol et al., 2011). However, just as knowing musical notes does not provide an understanding of music, and listing ingredients does not give a recipe for a dish, enumerating factors does not explain how exactly the complex, multi-layered, living system we call soil forms from lifeless rock. We still need to answer the key question: *how*? What are those internal driving forces, those "gears" and "mechanisms" that, interacting under the influence of external factors, create the astonishing diversity of soil cover on our planet?

The answer to this question is provided by the study of elementary soil-forming processes (ESFPs).

Definition of ESFPs

Elementary soil-forming processes (ESFPs) are the totality of specific physical, chemical, and biological reactions and phenomena that occur within the parent material and lead to the formation, transformation, and translocation of substances, resulting in the development of characteristic soil horizons and the profile as a whole (Birkeland, 1984; Scheffer et al., 2018).

Simply put, ESFPs are the specific "building blocks" from which any soil is constructed. It is the combination, intensity, and sequence of these processes that determines whether a Chernozem, a Podzol, or a humid tropical Ferralsol will form at a given site.

Four Fundamental Categories of Processes

The entire diversity of soil-forming processes can be conditionally divided into four basic categories, as proposed by Roy Simonson (Simonson, 1978):

1. Transformations — processes in which some substances within the soil are converted into others. Classic examples: weathering of primary minerals and formation of secondary clay minerals, decomposition and humification of organic residues.

2. Translocations — transfer of substances from one part of the soil profile to another. This movement can occur both downward with percolating water (eluviation) and upward (illuviation).

3. Additions (Accretions) — input of substances into the soil from outside. This includes: input of organic matter to the surface, deposition of atmospheric dust and aerosols, input of chemical elements with precipitation.

4. Losses (Eluviation) — removal of substances beyond the soil profile. This can include loss of gaseous compounds (e.g., carbon dioxide during respiration), removal of soluble salts and colloids with groundwater flow, or physical removal of material through erosion.

Any specific ESFP affects one or more of these basic categories to varying degrees. Understanding this principle is the foundation for a conscious perception of the entire subsequent classification of soils.

Processes and Factors: Two Sides of the Same Coin

It is important to emphasize that ESFPs serve as the link between soil-forming factors and the final properties of the soil. Factors create certain conditions (e.g., cold and humid climate), and under these conditions, specific processes begin to actively develop (e.g., podzolization). It is the combination of processes that forms the diagnostic horizons by which we distinguish soil types. As Birkeland (1984) writes, we are often forced to interpret processes "based on a combination of field and laboratory data."

Thus, a logical chain can be constructed:

Factors (Climate, Relief, Parent Material, Organisms, Time) → ESFPs (Humification, Podzolization, Lessivage, etc.) → Soil Horizons (A, E, B, C, etc.) → Soil Type (Sod-Podzolic, Chernozem, etc.)

Definition and Essence of the Process

Humification is the process of converting fresh organic residues (plant litter, root exudates, animal and microbial remains) into a complex, high-molecular-weight, stable complex of dark-colored compounds resistant to further decomposition, called humus (Foth, 1990; Scheffer et al., 2018).

This process is central to the formation of the humus-accumulative horizon (A), which determines soil fertility. Humification is not simple decomposition but a qualitative transformation: from amorphous plant mass, a new, colloidal, chemically active substance is born. Humus has an enormous specific surface area, high cation exchange capacity (CEC can reach 200–300 cmol(c)/kg), and the ability to bind soil particles into aggregates, which determines soil structure and water stability (Huang et al., 2012).

Two Sides of One Process: Mineralization and Humification

It is important to immediately draw a clear distinction between two parallel processes occurring in any organic material entering the soil:

• Mineralization — the complete decomposition of organic matter to inorganic compounds (CO₂, H₂O, NH₄⁺, NO₃⁻, PO₄³⁻, SO₄²⁻, etc.). This is the process of "burning" organic carbon by microorganisms to obtain energy. Mineralization is a loss of carbon from the soil.
• Humification — partial decomposition and subsequent synthesis, as a result of which carbon is retained in the soil but in a new, stable form. This is a process of carbon accumulation.

The ratio between these two processes determines whether the soil will accumulate organic matter or lose it. This ratio depends on a number of factors, primarily the conditions in which decomposition occurs (Foth, 1990; Weil, 2017).

Factors Determining the Intensity of Humification

Since humification is a biochemical process, all factors affecting the activity of microorganisms (the main "workers" of the soil) will determine its rate and direction.

1. Moisture and Aeration. Optimal conditions for humification are moderate moisture and good aeration (oxygen availability). In waterlogged, anaerobic conditions (e.g., in bogs), the decomposition process slows down sharply, and organic matter accumulates in the form of peat (not humus). In dry conditions, microbial activity is also suppressed (Scheffer et al., 2018; Weil, 2017).

2. Temperature. Chemical and biological reactions accelerate with increasing temperature. However, in the hot and humid climate of the tropics, the rate of mineralization is so high that humus does not have time to accumulate. The largest humus reserves are characteristic of temperate and subboreal climates, where sufficient moisture combines with seasonal cooling that slows mineralization (Foth, 1990; Weil, 2017). As the authors of "Scheffer/Schachtschabel" (2018) aptly note, "in warm regions, the cycle of matter proceeds faster, which does not always lead to humus accumulation."

3. Quality of Organic Material (Litter Composition). Easily decomposable compounds (sugars, proteins, cellulose) mineralize rapidly. Compounds resistant to decomposition—lignin, resins, waxes, tannins—are the main precursors of humic substances. Coniferous litter is rich in lignin and resins, poor in bases, and decomposes slowly, giving rise to coarse humus (mor). Deciduous litter and especially herbaceous vegetation contain more nitrogen and bases, decompose faster, and produce "mild" humus (mull) (Weil, 2017; Scheffer et al., 2018).

4. pH of the Environment. In acidic soils (pH < 5), bacterial numbers decline, and fungi begin to play a dominant role. Their metabolic products contribute to the conservation of organic residues and the formation of acidic forms of humus. In neutral and slightly alkaline environments, bacterial activity is maximal; decomposition and humification proceed more completely, forming fertile "mull" (Huang et al., 2012; Weil, 2017).

5. Texture and Mineralogy. Fine-grained minerals (clay particles) and iron and aluminum oxides have high adsorption capacity. They bind organic molecules and protect them from further mineralization. This is why humus predominantly accumulates in the fine fractions of soil (silt and clay). Soils with high clay content are generally richer in humus than sandy soils (Huang et al., 2012).

Morphological Expression: Diversity of Humus Forms

The result of humification is clearly manifested in the structure of the upper part of the soil profile. Different combinations of conditions form different humus forms (Foth, 1990; Scheffer et al., 2018):

• Mull: Favorable conditions (neutral environment, rich herbaceous vegetation, active fauna). Characterized by intensive mixing of organic matter with the mineral part of the soil. Features a thick, dark, structural A horizon (with signs of earthworm and other invertebrate activity). The most fertile form.
• Moder: An intermediate variant. Observed under slightly acidic conditions (e.g., under deciduous forests). Organic litter decomposes but does not completely mix with the mineral horizon. A distinct litter layer (O-horizons) and a well-expressed but less thick humus horizon (Ah) are formed.
• Mor: Extreme conditions (acidic environment, nutrient-poor coniferous litter, lack of soil fauna). Decomposition proceeds extremely slowly; organic material accumulates in the form of a thick, poorly decomposed litter layer (O-horizon). The transition to the mineral horizon is sharp. This form is characteristic of podzolic soils of the taiga zone.

Biological Basis: Who Creates Humus?

The main creators of humus are three groups of soil organisms (Scheffer et al., 2018; Eash et al., 2016):

1. Macrofauna (especially earthworms and millipedes): They mechanically shred plant litter, mix it with the mineral part of the soil, and pass it through the digestive tract, where it is mixed with enzymes. Their "coprolites" (casts) are the basis of water-stable granular structure in mull.

2. Microfauna (mites, springtails, collembola): They feed on bacteria and fungi, contributing to further processing of organic matter and release of nutrients.

3. Microorganisms (bacteria and, especially, actinomycetes and fungi): They play the key role in the chemical transformation of lignin and cellulose molecules into the complex aromatic structures of humic acids.

Practical Significance of Humification

The intensity and nature of humification determine the key agronomic properties of the upper horizon: its color (the darker, the more humus), reserves of nutrients (especially nitrogen and phosphorus), water-holding capacity, structure, and resistance to erosion. As Weil (2017) notes, soils with high humus content act as "islands of fertility," exerting a decisive influence on the productivity of natural and agroecosystems.

Thus, humification is a fundamental biogenic process of organic matter accumulation that creates the basis of soil fertility. However, along with humus accumulation, other processes—both constructive and destructive—constantly occur in the soil, to which we now turn.

Definition and Essence of the Process

Mineralization is the process of decomposing soil organic matter and plant residues to final inorganic products: carbon dioxide, water, ammonium, nitrates, phosphates, sulfates, and other mineral compounds (Foth, 1990; White, 2006).

If humification is a process of carbon conservation and creation of a stable organic reservoir, then mineralization is a process of releasing the chemical energy and nutrients stored in organic matter. It is carried out by the same microorganisms that participate in humification but is directed toward the opposite—the return of elements to the soil solution and the atmosphere.

As the authors of "Scheffer/Schachtschabel" (2018) figuratively formulate, mineralization is the "burning" of organic matter without flame, during which carbon in the form of CO₂ escapes into the atmosphere, and nitrogen, phosphorus, sulfur, and potassium become available for a new cycle of life.

Mineralization as an Integral Part of the Biological Cycle

In natural ecosystems, mineralization and humification occur simultaneously but in different proportions. It is the balance of these two processes that determines whether the soil will accumulate organic matter or lose it.

Simplified, this can be represented as a cycle:

Plants (photosynthesis) → Organic matter (litter, roots) → Microorganisms (humification + mineralization) → Inorganic elements (N, P, K, S) → Plants (uptake) → New organic litter → ...

Thus, mineralization closes the biological cycle of elements, ensuring continuous plant nutrition even without external fertilizers (Eash et al., 2016; Weil, 2017).

Two Types of Microorganisms Involved in Mineralization

Mineralization is carried out by two main groups of microorganisms, differing in their living conditions:

1. Aerobic microorganisms (bacteria, fungi, actinomycetes): They use atmospheric oxygen as the final electron acceptor in the oxidation of organic matter. The reaction proceeds according to the scheme:

Aerobic mineralization proceeds quickly and completely, with the formation of mainly carbon dioxide and water. This is the dominant process in well-drained, aerated soils (White, 2006).

2. Anaerobic microorganisms (mainly bacteria): In conditions of oxygen deficiency (waterlogging, compaction), microorganisms use other oxidized compounds as electron acceptors (nitrates, sulfates, iron and manganese oxides). This process proceeds more slowly and is accompanied by the formation of reduced gases—methane (CH₄), hydrogen sulfide (H₂S), ammonia (NH₃), and nitrous oxide (N₂O) (Huang et al., 2012; Kimura et al., 2012). Anaerobic mineralization is characteristic of bog, gley, and paddy soils.

Factors Determining the Rate of Mineralization

The intensity of mineralization is regulated by the same factors as humification, but the direction of their effect may be opposite:

• Temperature: With increasing temperature, the rate of microbial activity and, consequently, the rate of mineralization increases. The highest rate is achieved in warm, humid tropics, where organic matter mineralizes almost completely, and humus does not accumulate. In cold regions, mineralization is slowed, which promotes organic matter accumulation (Foth, 1990; Weil, 2017).
• Moisture: Optimal moisture (close to field capacity) stimulates mineralization. Under waterlogging (anaerobiosis), the process slows down and shifts to a reductive pathway. Under drying, microbial activity is also suppressed (White, 2006).
• pH: Neutral and slightly alkaline reaction (pH 6.5–7.5) is most favorable for bacterial mineralization. In acidic soils, fungi dominate, which decompose organic matter more slowly, although they are more effective at breaking down lignin (Weil, 2017).
• Composition of Organic Matter: Readily available compounds for microorganisms (sugars, proteins, amino acids) mineralize within days or weeks. Complex polymers (cellulose, hemicellulose)—over months. Lignin and humic substances are the most resistant; their mineralization lasts for years and even decades (Huang et al., 2012).
• Texture and Mineralogical Composition: Fine clay particles and iron and aluminum oxides adsorb organic molecules, making them less accessible to microbial enzymes, thereby slowing their mineralization. In sandy soils, lacking such "protection," mineralization proceeds faster (Birkeland, 1984; Huang et al., 2012).

Mineralization and Nutrient Availability

The main practical significance of mineralization lies in its role as a supplier of nutrients in plant-available form. It is during mineralization that:

• Organic nitrogen is converted to ammonium (NH₄⁺) and nitrate (NO₃⁻) forms.
• Organic phosphorus is released as phosphate ions (H₂PO₄⁻, HPO₄²⁻).
• Sulfur is released as sulfate ion (SO₄²⁻).

This process is called "biological release" or "immobilization-mineralization" (Eash et al., 2016; Foth, 1990). It is important to understand that if fresh organic matter with a wide C:N ratio (e.g., straw) is added to the soil, actively multiplying microorganisms may temporarily "immobilize" (bind into their biomass) available nitrogen, causing its temporary deficiency for plants. However, after the death of this microbial biomass, nitrogen is again mineralized and becomes available.

Ecological Significance: Mineralization and the Global Carbon Cycle

Mineralization is one of the main sources of CO₂ input into the atmosphere. Globally, approximately 10 times more carbon is released from soils annually (mainly through mineralization) than is burned from fossil fuels (Weil, 2017). Therefore, the balance between carbon sequestration through humification and its release through mineralization is of critical importance for the planet's climate system. Any change in conditions (e.g., warming or land-use change) that accelerates mineralization leads to additional greenhouse gas emissions.

Mineralization as Part of Soil Respiration

Mineralization of organic matter is the main component of soil respiration (CO₂ release). In addition to the actual decomposition of organic matter, respiration also includes plant root respiration. However, it is mineralization that determines the long-term dynamics of soil organic carbon. By measuring soil respiration, we indirectly assess the rate of mineralization and the activity of the entire microbial community (Gregorich et al., 2012).

Thus, mineralization is the reverse side of humification. Together, they form a dynamic equilibrium that maintains the constant cycling of carbon and nutrients in the biosphere. If humification creates a reserve of organic matter in the soil, then mineralization returns it to active biological turnover. Understanding this balance is key to managing fertility and sustainability of agroecosystems.

Definition and Essence of the Process

Lessivage (from French lessivage — "washing out," "leaching") is the process of mechanical translocation of clay particles from the upper horizons of the soil to the lower ones, with their subsequent accumulation in the illuvial horizon (Birkeland, 1984; White, 2006).

In Russian literature, this process is often called "illuvial-eluvial differentiation by clay" or simply "clay eluviation." However, lessivage is not simply clay transport but a complex three-phase mechanism involving:

1. Dispersion (peptization) of clay particles—their transition from an aggregated state to a suspension.

2. Transport of clay particles with the downward flow of soil moisture.

3. Precipitation (flocculation) of clay in the lower part of the profile with the formation of a horizon enriched in clay (Bt-horizon).

It is important to emphasize that during lessivage, the chemical composition of clay particles does not change—they are only physically translocated, unlike processes of podzolization or ferralitization, where dissolution and reprecipitation of substances with changes in their composition occur (White, 2006; Birkeland, 1984).

Conditions Necessary for Lessivage Development

Lessivage is a process requiring a specific combination of factors. It does not occur in any soil but manifests only under a number of conditions:

1. Presence of clay particles in the parent material or their formation through weathering. The parent material must contain a sufficient amount of fine-grained minerals capable of peptization (predominantly 2:1-layer silicates such as montmorillonite, hydromicas, vermiculite) (Foth, 1990; Buol et al., 2011).

2. Sufficient but not excessive moisture. Lessivage develops under leaching or periodically leaching water regimes when precipitation exceeds evapotranspiration but does not yet lead to permanent anaerobiosis. Climates with pronounced seasonal moisture redistribution (Mediterranean, continental temperate) are optimal (Birkeland, 1984; Scheffer et al., 2018).

3. Neutral or slightly acidic reaction (pH 5.5–7.0). At low pH (strongly acidic conditions), clay particles lose their negative charge due to adsorption of Al³⁺ cations and coagulate, becoming immobile. At high pH (alkaline conditions), clays may disperse, but this leads to solonetzation (see Section 8), not classical lessivage (White, 2006; Weil, 2017).

4. Presence of exchangeable cations promoting dispersion. Monovalent cations (Na⁺, K⁺) or divalent cations at low concentration (Ca²⁺, Mg²⁺) most effectively disperse clay. However, with excess Ca²⁺ (especially in carbonate soils), clays flocculate, and lessivage does not occur. This is why lessivage develops most actively after leaching of carbonates from the upper part of the profile (Birkeland, 1984; Scheffer et al., 2018).

5. Alternation of wetting and drying periods. Upon drying, soil cracks and macropores form, through which water with clay suspension can penetrate deeply during subsequent wetting. Therefore, lessivage is particularly intense in climates with a pronounced dry season or in soils with seasonal freezing-thawing (Foth, 1990; Weil, 2017).

Mechanism of the Process: Three Stages

Morphological Features of Lessivage

Lessivage leaves characteristic diagnostic features in the soil profile by which it can be recognized in the field:

1. Eluvial Horizon (E or A-E): The upper part of the profile from which clay has been removed. It is discolored (light gray, pale), often has platy or foliated structure, and is depleted in the clay fraction (Foth, 1990; Weil, 2017).

2. Illuvial Clay Horizon (Bt): The underlying horizon enriched in clay. It has a darker (brown, reddish-brown) color, is denser, often with nutty or prismatic structure. Shiny clay films—cutans—are visible on the faces of structural peds (Birkeland, 1984; White, 2006).

3. Textural Differentiation: A clear increase in clay content from the upper horizons to Bt. For diagnosing lessivage in classification systems, a quantitative criterion is often used: clay content in Bt must be at least 3–4% (absolute) higher than in the eluvial horizon (Buol et al., 2011).

4. Characteristic Transition: The transition between the eluvial and illuvial horizons can be sharp or gradual, sometimes tongue-shaped (clay material trickling down cracks).

Difference of Lessivage from Other Translocation Processes

Lessivage is often confused with other processes of substance transport. However, there are fundamental differences between them:

(White, 2006; Birkeland, 1984; Weil, 2017)

Relationship with Other Processes: Sequence and Combination

Lessivage rarely acts in isolation. In nature, it is often combined with or replaced by other ESFPs:

• At initial stages (in young, not yet oxidized soils), lessivage may be the leading process, forming so-called "lessivated" or "podzolized" soils (Buol et al., 2011).

• With increasing acidity (e.g., under coniferous forest influence), lessivage may transition to podzolization—clay particles in an acidic environment are destroyed, and removal begins not of clay but of its breakdown products—Fe, Al, and silica (White, 2006; Birkeland, 1984).

• In alkaline, soda conditions (salinization), lessivage transforms into solonetzation, where sodium promotes clay dispersion and formation of a dense columnar horizon (Scheffer et al., 2018).

Thus, lessivage is a process characteristic of transitional conditions between humid and arid climates, and its presence in the profile often indicates past or present periods of moderate leaching.

Ecological and Agronomic Significance of Lessivage

Lessivage has a dual effect on soil fertility:

Positive aspects:

• Forms a clay illuvial horizon (Bt), which has high cation exchange capacity and serves as a kind of "accumulator" of nutrients, preventing their leaching beyond the root zone (Foth, 1990).
• Improves aggregation and structure both in the eluvial horizon (due to clay removal—it becomes looser, more air-permeable) and in the illuvial part of the profile (due to the cementing action of clay films) (Weil, 2017).

Negative aspects:

• Formation of a dense Bt-horizon with low hydraulic conductivity can lead to periodic waterlogging of the upper part of the profile (formation of perched water, development of gleying). This is especially noticeable in soils with a sharp textural break (Foth, 1990; Scheffer et al., 2018).
• Depletion of the upper horizon in clay reduces its cation exchange capacity and water-holding capacity, making it more susceptible to erosion and drying.
• Lessivated soils often require special tillage practices (deep tillage, application of organic matter to improve structure).

Geographic Distribution

Lessivage is most widely distributed in soils of:

• Temperate zone—brown forest soils (Cambisols), gray forest soils, some sod-podzolic soils (with elements of lessivage).
• Subtropical and Mediterranean regions—brown and red soils with well-developed Bt-horizon (Buol et al., 2011).
• Continental regions with seasonal precipitation distribution—forest-steppe and steppe zones, where Chernozems with signs of lessivage occur (especially with shallow carbonate horizon) (Scheffer et al., 2018).

In soil classification, lessivage is a key process for the identification of argic horizons (in Soil Taxonomy) and Luvisols, Retisols, Acrisols, Alisols (in WRB)—soils in which textural differentiation is a diagnostic feature (Buol et al., 2011; IUSS Working Group WRB, 2014).

Definition and Essence of the Process

Podzolization (or podzolic process) is the process of destruction of primary and secondary minerals in the upper part of the soil profile under the action of acidic organic substances, followed by removal of the breakdown products (primarily iron, aluminum, and organic matter) to the underlying horizons, where they accumulate (Birkeland, 1984; White, 2006).

The term "podzol" comes from the Russian word "зола" (zola) —"ash"—and indeed, the upper, bleached horizon of these soils resembles the color of wood ash. Podzolization is one of the most intense destructive processes in soil science, as a result of which the upper part of the profile loses not only fertility but also the very ability to retain nutrients (Weil, 2017).

Unlike lessivage, where clay particles are transported mechanically and remain chemically unchanged, during podzolization, chemical destruction of minerals occurs with the formation of mobile organo-mineral complexes that migrate downward through the profile (White, 2006; Scheffer et al., 2018).

Chemical Basis of the Process: The Role of Organic Acids

The key feature of podzolization is complexation (or chelation). In the upper part of the profile (in the litter and humus horizon), during the decomposition of coniferous litter, heather, and other plants, fulvic acids are formed—low-molecular-weight organic acids with high complexing ability (Huang et al., 2012; White, 2006).

These acids, possessing numerous carboxyl and hydroxyl groups, interact with metal cations (especially Fe³⁺ and Al³⁺), forming soluble metal-organic complexes (chelates). Schematically:

The formed chelates readily dissolve in water and migrate downward through the profile. This process is called "cheluviation" (from chel + eluviation) (White, 2006; Birkeland, 1984).

Main Stages of Podzolization

Morphological Features of Podzolization

As a result of podzolization, a characteristic contrasting profile is formed, clearly distinguishable in the field (Foth, 1990; Weil, 2017):

1. Litter (O-horizon): Thick, acidic, often consisting of poorly decomposed coniferous needles and mosses. Subdivided into L (fresh litter), F (fermentation—partially decomposed), and H (humified—fully decomposed but not mixed with the mineral part) (Scheffer et al., 2018).

2. Eluvial Horizon (E or A₂): Light gray, ashy, or whitish layer. This horizon gave the process its name (from "ash"). It has almost completely lost not only Fe and Al but also all fine earth—only weathering-resistant minerals (quartz) remain. The horizon has a structureless or weakly developed platy appearance and is very poor in nutrients (White, 2006).

3. Illuvial Horizon (B): Dark brown, ochre-reddish, or almost black horizon where accumulated substances are deposited. Depending on the predominance of a particular component, subhorizons may be distinguished:

• Bh—predominantly humus accumulation (organic complexes).
• Bs—predominantly accumulation of iron and aluminum oxides (Fe, Al).
• Bhs—mixed accumulation (humus + Fe, Al) (Birkeland, 1984; Scheffer et al., 2018).

4. Transitional Horizon (BC or C): Disturbed or weakly weathered material.

The most pronounced forms of podzolization are characterized by sharp, sometimes "tongue-shaped" (invading) boundaries between horizons, reflecting the depth of penetration of organic acids along cracks and macropores (White, 2006).

Conditions Necessary for Podzolization

Podzolization is a strictly ecologically conditioned process. Its active development requires a combination of several factors (Foth, 1990; Birkeland, 1984; Weil, 2017):

1. Acidic, base-depleted parent material. The process proceeds most intensively on quartz sands, granites, acidic shales, and other rocks with low carbonate content and lacking alkaline and alkaline-earth elements capable of neutralizing acids.

2. Acidic, hard-to-decompose vegetation. Coniferous forests (spruce, pine, fir, larch), as well as heathlands (Calluna, Vaccinium), produce litter rich in lignin, tannins, and resins, which upon decomposition yield strong organic acids (fulvic acids) and slow the activity of decomposer microorganisms (Scheffer et al., 2018).

3. Cool, humid climate (high precipitation:evapotranspiration ratio > 1). It is in the taiga and tundra zones (boreal belt) that optimal conditions for podzolization are created: excess moisture promotes profile leaching, and low temperatures slow organic matter decomposition and promote accumulation of acidic products (Foth, 1990; Weil, 2017).

4. Good drainage (free water removal). Water stagnation and anaerobiosis hinder the formation of organic acids and podzolization (in waterlogged conditions, gleying develops, which will be discussed in the next section) (Birkeland, 1984).

Difference of Podzolization from Lessivage

Since both processes lead to the appearance of a light upper and darker lower horizon, they are often confused. However, there is a fundamental difference between them:

(White, 2006; Birkeland, 1984; Weil, 2017)

Genetic Relationship with Lessivage: Bisolums

In nature, podzolization is often superimposed on lessivage. This phenomenon is called "bisolums" —soils with two successive eluvial-illuvial cycles (Scheffer et al., 2018; Birkeland, 1984).

For example, in sod-podzolic soils of central Russia, a "double" profile is often observed:

1. Upper cycle—modern podzolization (E + Bh, Bs).

2. Lower, deeper cycle—relict lessivage (lower eluvial horizon above Bt)—a remnant from previous, warmer climatic epochs when conditions were less acidic (White, 2006).

Thus, the modern profile of a sod-podzolic soil is the result of the superposition of two different ESFPs over time.

Evolution of Podzols Over Time: Stages of Maturity

Podzolization is a long-term process that passes through a series of stages (Birkeland, 1984; Foth, 1990):

1. Initial stage (weak podzolization): Only lightening of the upper part of the profile appears (patchy, tongue-shaped). Iron and aluminum have not yet reached deep horizons. Only a weak Bs-horizon may be expressed in the form of light ochreous coatings.

2. Developed stage (moderate podzolization): A distinct E-horizon 10–15 cm thick forms, and individual lenses of Bhs appear in the lower part of the profile. The transition between horizons is not yet sharp.

3. Mature stage (strong podzolization, Podzol): E-horizon is thick (up to 30–40 cm), sharply contrasting with dark brown Bhs or black Bs-horizon. Orstein (cemented layers of Fe-humus compounds) occurs, which are practically impenetrable to roots and water (White, 2006; Scheffer et al., 2018).

Biological Role in Podzolization: "Fungal Decomposition"

The important role of mycorrhizal fungi in podzolization should be noted separately. Modern research shows that some fungi living in symbiosis with tree roots can "leach" minerals directly, extracting iron and other elements from them bypassing acid dissolution (van Breemen et al., 2000; Jongmans et al., 1997, cited in White, 2006). Fungal hyphae penetrate into microfractures of minerals and release organic acids locally, sharply accelerating destruction. This process is called "mineral tunneling" (Barker et al., 1997, cited in White, 2006).

Geographic Distribution of Podzols

Podzols and podzolic soils are among the most widespread soil types on Earth. They occupy vast areas:

• Boreal belt (taiga)—Scandinavia, Northern Europe (Finland, Sweden, Norway), Russia (from the Kola Peninsula to the Far East), Canada, Northern USA (Alaska, northern states).
• High mountains and coastal areas with humid cool climates—Alps, Carpathians, Western Cordilleras of North and South America, New Zealand, Tasmania (Buol et al., 2011; IUSS Working Group WRB, 2014).
• Extratropical heathlands and moorlands—e.g., in England, Germany (North German Plain) (Scheffer et al., 2018).

In soil classification systems, podzolization is a diagnostic process for the identification of Spodosols (Soil Taxonomy) and Podzols (WRB) (Buol et al., 2011; IUSS Working Group WRB, 2014).

Ecological and Agronomic Significance

Podzolization leads to a sharp decline in natural soil fertility:

• Eluvial horizon (E) loses almost all nutrients (Ca, Mg, K, Fe, Mn, Zn, Cu). Only quartz sand remains, incapable of retaining water and nutrients (Foth, 1990; Weil, 2017).
• Illuvial horizon (Bhs) concentrates those elements that were removed from E, but in forms often poorly available to plants (Fe and Al compounds strongly bind phosphorus, making it unavailable).
• Acidic reaction throughout the profile (pH 3.5–4.5) suppresses the activity of most beneficial microorganisms.

However, with cultivation (liming + fertilization), podzolic soils can be significantly improved. Many agricultural crops (potatoes, rye, flax, perennial grasses) are successfully grown on them, provided high doses of organic and mineral fertilizers and liming are applied (Scheffer et al., 2018; Eash et al., 2016).

Podzolization in the Past and Present

It is important to understand that many Podzols observed today did not form in modern conditions but in past cold periods (Pleistocene) when the climate was harsher. Often podzolic soils are polygenetic—relict formations that have survived to the present day, but the process of podzolization in them may currently be weakened or even stopped (Birkeland, 1984). This once again emphasizes the importance of considering the time factor when interpreting soil profiles.

Definition and Essence of the Process

Gleying (or gleyzation) is the process of soil transformation under conditions of prolonged or periodic waterlogging, accompanied by oxygen deficiency (anaerobiosis) and the development of reducing conditions, leading to changes in soil color, redistribution, and partial removal of iron and manganese (Birkeland, 1984; White, 2006).

The term comes from the Russian word "глей" (gley) —the name given to the viscous, structureless, bluish or bluish-gray clayey mass that forms in the lower part of soils under conditions of constant excessive moisture. This is one of the most striking and morphologically expressed processes in soil science, as it fundamentally changes the color of the profile, making it gray, bluish, or greenish (Foth, 1990; Weil, 2017).

Gleyzation is a process that manifests to varying degrees in soils of all climatic zones—from tundra to humid tropics, wherever conditions for prolonged water stagnation and oxygen deficiency are created. As White (2006) notes, "gleying is the most widespread process in hydromorphic (waterlogged) soils of the world."

Chemical Basis of the Process: Reduction of Iron and Manganese

The key chemical difference of gleying from most other ESFPs is reduction reactions (redox transformations) occurring under anaerobic conditions. In normal, aerated soils, iron and manganese are in oxidized form: iron as Fe³⁺ (insoluble oxides and hydroxides, coloring the soil yellow, orange, red, and brown tones), manganese as Mn⁴⁺ (black oxides).

Under oxygen-deficient conditions, microorganisms performing anaerobic respiration are forced to use not oxygen but other compounds as electron acceptors, including iron and manganese oxides. As a result (Birkeland, 1984; White, 2006; Weil, 2017):

The formed divalent ions Fe²⁺ and Mn²⁺ become mobile—they pass into the soil solution and can migrate both within the profile and beyond it. It is the disappearance or sharp decrease in the content of oxidized forms of iron and manganese that leads to the appearance of the bluish, greenish-gray, or bluish-gray color characteristic of gleying (Foth, 1990; Scheffer et al., 2018). As the authors of Scheffer/Schachtschabel (2018) figuratively put it, "gleying is the process of 'washing out' color, when the soil loses its familiar reddish and brown tones."

The Role of Microorganisms in Gleying

Gleying is not just a physicochemical but a biogeochemical process. Its initiation and maintenance are possible only with the active participation of reducing microorganisms (Kimura et al., 2012; Huang et al., 2012):

1. Facultative and obligate anaerobes—bacteria capable of living and reproducing in an oxygen-free environment. These include many representatives of the genera Pseudomonas, Bacillus, Clostridium, Desulfovibrio, and others.

2. Sulfate-reducing bacteria—reduce sulfates to sulfides, which, in the presence of Fe²⁺, leads to the formation of black iron sulfides (FeS, FeS₂)—the so-called "agricultural hydrogen sulfide" (Kimura et al., 2012).

3. Iron-reducing bacteria (e.g., Geobacter, Shewanella)—directly use Fe³⁺ as an electron acceptor, converting it to Fe²⁺ (Huang et al., 2012).

It is important to understand: without organic matter (an energy source for microorganisms), gleying does not occur even under flooding conditions. This is why gleying develops most intensively in soils rich in easily decomposable organic matter (meadows, bogs, rice paddies) (Kimura et al., 2012; Weil, 2017).

Conditions Necessary for Gleying Development

For active gleying to occur, a combination of three main factors is necessary (Birkeland, 1984; Foth, 1990; White, 2006):

1. Excessive moisture (water stagnation). Water fills the pore space, hindering oxygen diffusion from the atmosphere. The rate of oxygen consumption by microorganisms in this case exceeds the rate of its supply, and anaerobic conditions are created. This can be caused by:

• High groundwater table (groundwater gleying).
• Presence of a water-restricting horizon (e.g., clayey Bt or orstein), creating perched water (surface gleying).
• Prolonged flooding (floodplains, rice paddies).

2. Presence of organic matter. It serves as an energy source for microorganisms and an electron donor for the reduction of Fe³⁺ and Mn⁴⁺. The more organic matter and the more readily it decomposes, the more intensive the gleying.

3. Sufficient temperature for microbial activity. In cold climates, gleying also develops but more slowly (e.g., in tundra gley soils). Optimal temperatures for active gleying are within 15–35 °C (Kimura et al., 2012; Weil, 2017).

Morphological Features of Gleying

Gleying leaves characteristic, easily diagnosable features in the soil profile (Foth, 1990; Weil, 2017; Scheffer et al., 2018):

1. Uniform reduced coloration (reductomorphic): The soil acquires bluish, bluish-gray, greenish, sometimes almost white tones. This coloration is due to the presence of ferrous iron (Fe²⁺) and manganese (Mn²⁺) compounds. Such horizons are designated by the symbol G (from gley)—in German, international, and American nomenclature (Birkeland, 1984; White, 2006).

2. Rusty-ochreous and black spots (oxidized forms): During periodic drying or near plant roots (where oxygen is accessible), Fe²⁺ oxidizes to Fe³⁺, precipitating as bright rusty, yellow-brown spots and concretions. Manganese upon oxidation produces black dotted inclusions and films. Such "mottled" coloration is characteristic of semi-gley and seasonally waterlogged soils. In international nomenclature, the symbol g is used to designate such horizons (in combination with other letters—e.g., Bg, Cg) (White, 2006).

3. Signs of iron reduction in field tests: If a drop of α,α-dipyridyl solution (a reagent for Fe²⁺) is applied to such soil, it turns red—a reliable field indicator of reduced iron (Weil, 2017).

4. Characteristic structure: Gley horizons are often structureless, viscous (silty), cracking into large clods upon drying, but having a "soapy" consistency in the wet state (Foth, 1990; Scheffer et al., 2018).

Types of Gleying: Groundwater and Surface

Depending on the cause of waterlogging, two main types of gleying are distinguished (White, 2006; Birkeland, 1984; Scheffer et al., 2018):

Gleying as a Secondary Process

Gleying often develops as a secondary, superimposed process on already formed soils. For example:

• Lessivated soils with a dense Bt-horizon, with increased moisture, can become surface-gleyed (pseudogleys).
• Podzols with an orstein (cemented) B-horizon also create perched water and can gley above the water-restricting layer (Birkeland, 1984; White, 2006).
• Even Chernozems with prolonged flooding (e.g., after dam construction or drainage disruption) can acquire signs of gleying in the lower part of the profile (Foth, 1990).

Thus, gleying is not only an independent process but also an important indicator of changing moisture conditions both in the modern period and in the past (paleogleying).

Sulfidic Gleying (Special Case)

In some waterlogged soils, especially in the presence of sulfates (marine deposits, saline groundwater) and abundant organic matter, sulfidic gleying develops. Under anaerobic conditions, sulfate-reducing bacteria reduce sulfates to sulfides (H₂S), which interact with Fe²⁺ to form black iron sulfides—pyrite (FeS₂) and melnikovite (FeS) (White, 2006; Huang et al., 2012).

Such soils are called sulfidic or potential acid sulfate soils (in literature—sulfidic materials). Upon drainage and aeration, sulfides oxidize, forming sulfuric acid, leading to sharp acidification (pH < 4) and the formation of so-called acid sulfate soils—among the most aggressive and infertile soils in the world (White, 2006; Weil, 2017). This process will be discussed in more detail in Section 7. "Salinization."

Difference of Gleying from Other Processes

Gleying is often confused with podzolization (both produce light upper horizons) and with lessivage (both can form dense lower horizons). However, there are clear diagnostic criteria:

(White, 2006; Birkeland, 1984; Weil, 2017)

Geographic Distribution and Classification

Gleyed soils are distributed throughout the globe but especially widely:

• In boreal and subarctic regions (tundra, forest-tundra, taiga)—where permafrost and permanent waterlogging are widespread. These are cryogley soils.
• In the temperate zone with excessive moisture (Western Europe, northeastern USA, Eastern Europe)—here pseudogleys and gley soils form.
• On floodplains and river deltas of all climatic zones—floodplain gley soils.
• In the tropics and subtropics—rice soils (paddy), mangrove and swamp soils (White, 2006; Buol et al., 2011).

In classification systems, gleying is a diagnostic feature for the identification of Gleysols (WRB), Aquic suborders (Soil Taxonomy), as well as many soils with water regimes (Histosols, Fluvisols, Planosols, Stagnosols) (Buol et al., 2011; IUSS Working Group WRB, 2014).

Ecological and Agronomic Significance of Gleying

Gleying has a decisive influence on fertility:

Negative consequences:

• Oxygen deficiency in the root zone—roots of most cultivated plants cannot withstand prolonged flooding (exception—rice and some wetland plants).
• Accumulation of toxic reduced compounds—Fe²⁺, Mn²⁺, H₂S, organic acids—in concentrations that inhibit plants (Foth, 1990; Weil, 2017).
• Slowing of organic matter mineralization (due to anaerobiosis)—nutrient release is slowed, although organic matter accumulation in the form of peat increases (in extreme cases) (Scheffer et al., 2018).
• Low bearing capacity—gley soils often have weak bearing capacity and are unsuitable for construction without drainage.

Positive aspects (with managed use):

• Rice is successfully cultivated on gleyed soils—a crop adapted to prolonged flooding (Eash et al., 2016; Weil, 2017).
• With drainage (reclamation), such soils can become highly productive, especially with application of organic fertilizers and liming (for acidic gley soils) (Scheffer et al., 2018).

Gleying as a Paleoindicator

Since gleying is associated with water regime, the presence of gley horizons (especially relict ones with signs of modern drying) is an important paleoclimatic and paleogeomorphological indicator (Birkeland, 1984; Scheffer et al., 2018). From traces of ancient gleying, one can reconstruct:

• Former groundwater level.
• Former hydrological conditions (e.g., presence of lakes, bogs, oxbow lakes).
• Climatic changes (by the shift from gleying to oxidizing processes).

Thus, gleying is not only a modern process but also a "stone chronicle" of past hydrological conditions.

Definition and Essence of the Process

Salinization is the process of accumulation in the soil profile of water-soluble salts (chlorides, sulfates, carbonates, bicarbonates of sodium, calcium, magnesium, potassium) in concentrations that inhibit plant growth and deteriorate soil physical properties (Birkeland, 1984; Weil, 2017).

Salinization is one of the most widespread and economically significant degradation processes in soils of arid and semi-arid regions, where evapotranspiration significantly exceeds precipitation (Foth, 1990; White, 2006). According to estimates, more than 900 million hectares of soils are salinized worldwide, accounting for about 7% of all land and almost 25% of irrigated lands (Weil, 2017; Eash et al., 2016).

Salinization is the result of salt accumulation that enters the soil from various sources but cannot be leached due to insufficient moisture. As the authors of "Scheffer/Schachtschabel" (2018) figuratively formulate, "in dry regions, the cycle of salts is like a pendulum: they come but do not leave."

Sources of Salts Entering the Soil

Water-soluble salts enter the soil from several main sources (Weil, 2017; Birkeland, 1984; White, 2006):

1. Mineral composition of parent material. Some rocks (marine deposits, evaporite formations, saline continental deposits) are initially rich in readily soluble salts (NaCl, Na₂SO₄, CaSO₄·2H₂O, MgCl₂). Upon weathering of these rocks, salts are released and accumulate in the soil.

2. Atmospheric precipitation (cyclic salts). Winds blowing over the ocean lift tiny droplets of seawater into the atmosphere, which then fall with precipitation onto land. These are so-called "cyclic salts" —mainly NaCl and MgCl₂. In coastal areas, their contribution can reach significant values (up to 100–200 kg/ha/year) (White, 2006; Weil, 2017).

3. Dust accumulation. In desert and semi-desert regions, strong winds transport fine-grained material enriched in salts from one territory to another. This process is especially significant for soils of the Sahara, Gobi, Central Asia, and Australia (Birkeland, 1984).

4. Groundwater and hydrological regime. Saline groundwater rising through capillaries to the surface evaporates, leaving salts in the soil. This mechanism plays a key role in the development of salinization in closed depressions and in irrigated areas (Foth, 1990; Weil, 2017).

5. Anthropogenic sources. Irrigation with saline waters, improper drainage, application of mineral fertilizers in high doses, industrial emissions (especially sulfates), and even winter road treatment with de-icing mixtures (NaCl, CaCl₂, MgCl₂) can cause secondary salinization (Weil, 2017; Eash et al., 2016).

Conditions Necessary for Salinization Development

Salinization occurs with the simultaneous action of three main factors (Birkeland, 1984; White, 2006; Weil, 2017):

1. Arid or semi-arid climate with negative water balance. Precipitation is significantly less than evapotranspiration, so downward water flows cannot leach salts beyond the root zone.

2. Salt input. If salts do not enter, salinization does not develop even in dry conditions. The source can be parent material, groundwater, or the atmosphere.

3. Capillary rise of saline water. Water with dissolved salts rises through capillaries from groundwater or from deeper wetter horizons to the surface, where it evaporates, leaving salts. This mechanism is especially effective in fine-textured (loamy, clayey) soils where capillary rise is most pronounced (Foth, 1990; White, 2006).

Mechanisms of Salt Accumulation: Two Main Pathways

Depending on the source of salts and the nature of the water regime, two main types of salt accumulation are distinguished (Weil, 2017; Scheffer et al., 2018):

Morphological Features of Salinization

Salinized soils have characteristic, easily diagnosable features in the field (Foth, 1990; Weil, 2017; White, 2006):

1. Salt crust or efflorescences on the surface. At high salt concentrations, a loose white or grayish crust (scab) of NaCl, Na₂SO₄, CaSO₄·2H₂O crystals forms on the surface. This is a classic feature—Solonchak.

2. White coatings and crystals in pores and cracks. In the profile, small white inclusions—salt crystals filling pores and cracks, especially in the upper part of the profile—are visible.

3. Rusty and brown spots. In the presence of sulfates (especially FeSO₄), yellowish-brown, rusty spots may appear. In the case of sulfate salinization, the soil often has a yellowish-gray tint.

4. Characteristic structure. The upper salinized horizons are often loose, structureless, or weakly aggregated (powdery, coarse-grained). The lower horizons under hydromorphic salinization may be dense, sometimes with signs of gleying (bluish tints) (White, 2006).

5. Electrical conductivity (EC) as a diagnostic indicator. In laboratory conditions, salinization is determined by the electrical conductivity of a water extract (ECe). The salinization threshold is ECe > 4 dS/m (Weil, 2017; Eash et al., 2016). For strongly salinized soils, ECe values can reach 50–100 dS/m or more.

Classification of Salinization by Salt Composition

Depending on the predominant anions and cations, the following chemical types of salinization are distinguished (White, 2006; Weil, 2017; Birkeland, 1984):

1. Chloride salinization (dominated by NaCl, CaCl₂, MgCl₂). The most toxic to plants, as Cl⁻ in high concentrations causes osmotic shock and specific ion toxicity. Characteristic of coastal areas and some inland salt flats.

2. Sulfate salinization (dominated by CaSO₄·2H₂O, Na₂SO₄, MgSO₄). Less toxic than chloride but also inhibits plant growth at high concentrations. Characteristic of continental arid regions with gypsum-bearing rocks.

3. Carbonate and bicarbonate salinization (dominated by Na₂CO₃, NaHCO₃). Leads to alkaline salinization (pH > 8.5) and solonetzation. These salts are especially dangerous because alkaline environments are toxic to roots and promote clay dispersion (see next section—solonetzation).

4. Mixed salinization (most common in nature). Usually a mixture of chlorides, sulfates, and bicarbonates with predominance of one or another component depending on specific conditions.

Secondary Salinization (Anthropogenic)

This is salinization caused by human activity (Weil, 2017; Eash et al., 2016; Birkeland, 1984). The most common causes:

• Improper irrigation without drainage: Water evaporates, and salts brought with irrigation water accumulate in the upper horizon. Particularly dangerous is irrigation with waters of high mineralization (ECw > 1.5–2 dS/m) (Weil, 2017).
• Rising groundwater level: Without good drainage, groundwater level rises, and salts from deep horizons are drawn to the surface by capillary action.
• Disruption of natural water balance: For example, during reservoir creation, dam construction, swamp drainage, which changes the direction of saline water flows (Foth, 1990).

Secondary salinization is a serious problem in many irrigated regions of the world (Central Asia, China, India, Pakistan, USA, Australia). According to some estimates, up to 1–2 million hectares of irrigated land are lost annually due to secondary salinization (Weil, 2017).

Patterns of Salt Distribution in the Profile

Depending on the water regime and type of salinization, salts are distributed differently in the soil profile (Birkeland, 1984; White, 2006; Weil, 2017):

• In automorphic (dry) soils: Salt maximum is usually in the upper horizon (0–20 cm) or in the middle part of the profile, where maximum evaporation occurs. Below, salinization decreases.

• In hydromorphic (groundwater-capillary) soils: Maximum is often on the surface (salt crust), then a sharp decrease, but salts may occur throughout the profile, especially if groundwater is highly mineralized.

• Under irrigation: Characterized by salt accumulation in the upper 10–30 cm, especially in inter-rows where evaporation is maximal. Under irrigation furrows—a zone of desalinization (Foth, 1990; Eash et al., 2016).

Effect of Salinization on Plants

Salinization has a complex negative impact on plants (Eash et al., 2016; Weil, 2017):

1. Osmotic effect: Salts in the soil solution create high osmotic pressure that hinders water uptake by roots. The plant expends additional energy to maintain water balance.

2. Toxic effect of individual ions: Particularly dangerous are Na⁺, Cl⁻, HCO₃⁻ (toxic to many crops), as well as B (boron), Se (selenium), which may be present in toxic concentrations in salinized soils.

3. Disruption of mineral nutrition: Excess Na⁺ competes with K⁺ and Ca²⁺ for root uptake, causing potassium and calcium deficiency.

4. Deterioration of soil physical properties: During solonetzation (see next section), clay dispersion leads to compaction, crust formation, deterioration of aeration and infiltration.

Plant sensitivity to salinization varies. Most sensitive (ECe < 2 dS/m)—beans, peas, lettuce, strawberries. Most tolerant (ECe > 10 dS/m)—barley, cotton, sugar beet, some sorghums (Weil, 2017; Eash et al., 2016).

Reclamation of Salinized Soils

Main methods of combating salinization (Eash et al., 2016; Weil, 2017; Birkeland, 1984):

1. Leaching—repeated abundant irrigation (up to 2–3 thousand m³/ha) followed by removal of saline water through drainage network. Used with good drainage and sufficient water availability.

2. Drainage—creation of drainage systems (closed, open drainage, vertical drainage) for removal of saline groundwater and prevention of its capillary rise.

3. Gypsum application (CaSO₄·2H₂O)—to replace exchangeable Na⁺ with Ca²⁺ in soda salinization and solonetzation. Gypsum promotes displacement of Na⁺ and its leaching as neutral salt (Na₂SO₄) (White, 2006; Weil, 2017).

4. Bioreclamation—cultivation of salt-tolerant crops and halophytes that absorb salts and can be removed from fields, gradually reducing salinization.

5. Agrotechnical practices—field leveling, alternating irrigation with leaching, use of irrigation water with low mineralization, retention of plant residues on the surface (mulching) to reduce evaporation.

Relationship with Other ESFPs: Salinization, Solonetzation, and Solodization

Salinization often leads to the development of solonetzation (with Na⁺ and alkaline salt accumulation) and solodization (with subsequent desalinization, leaching of salts, and acid degradation). These processes are closely related and are often considered as a genetic series:

Solonchak (salinization) → Solonetzation → Solodization

This transition reflects the evolution of soils under changing water regimes (e.g., with increasing moisture) or during reclamation measures (Birkeland, 1984; White, 2006). Solonetzation will be the subject of the next section of our lecture.

Geographic Distribution of Salinized Soils

Salinized soils are widely distributed (Weil, 2017; White, 2006; Birkeland, 1984):

• Desert and semi-desert zone: Africa (Sahel, Namib Desert), Middle East (Arabian Peninsula, Iran, Iraq), Central Asia (Karakum, Kyzylkum, deserts of China), Australia, USA (Great Basin, California), South America (Atacama, Patagonia).
• Dry steppe and forest-steppe zone: Southern Russia, Ukraine, Kazakhstan, Mongolia, northern China.
• Irrigation zone of all continents (secondary salinization).
• Coastal lowlands and salt marshes (tidal zone, mangroves).

In soil classification systems, salinization is a diagnostic feature for the identification of Solonchaks (WRB), Salids (Soil Taxonomy), as well as many soils with saline horizons (Aridisols, Gypsisols, Calcisols) (Buol et al., 2011; IUSS Working Group WRB, 2014).

Definition and Essence of the Process

Solonetzation is the process of accumulation of exchangeable sodium (Na⁺) in the soil adsorption complex, accompanied by deterioration of soil physical properties (dispersion of clay particles, destruction of aggregates, formation of a dense structureless horizon) and increase in alkalinity (pH > 8.5) (Birkeland, 1984; White, 2006).

The term comes from the Russian name for these soils—"солонцы" (solonets). This is one of the most insidious and difficult-to-reclaim processes, as it affects not only chemical composition but also fundamentally changes the physical structure of the soil, making it practically impermeable to water, air, and roots (Weil, 2017; Scheffer et al., 2018).

Solonetzation is closely related to salinization but is the next stage of evolution of salinized soils under changing moisture conditions. If salinization is the accumulation of readily soluble salts (chlorides, sulfates, carbonates of sodium, calcium, magnesium), then solonetzation is the replacement of calcium and magnesium cations by sodium in the exchange complex after the removal (leaching) of neutral salts (White, 2006; Birkeland, 1984).

Difference of Solonetzation from Salinization

It is important to immediately draw a clear distinction between these two processes, as they are often mentioned together but have different natures and different consequences:

(Weil, 2017; White, 2006; Birkeland, 1984)

Chemical Basis: Mechanism of Solonetzation

Solonetzation develops in several stages (Birkeland, 1984; White, 2006; Weil, 2017):

1. Primary soda-type salinization: Salts containing Na⁺ and carbonate or bicarbonate ions (Na₂CO₃, NaHCO₃) enter or form in the soil. This can occur during weathering of sodium-containing silicates (e.g., plagioclases), during chemical interaction of neutral sodium salts with CaCO₃ (with formation of Na₂CO₃), or with soda entering from groundwater.

2. Na⁺ adsorption by soil colloids: In the presence of Na⁺ in solution (especially in the presence of carbonates), replacement of Ca²⁺ and Mg²⁺ by Na⁺ occurs on the surface of clay particles and humic substances:

This reaction is reversible, but at high Na⁺ concentration in solution (especially at pH > 8.5), it shifts to the right. The proportion of exchangeable Na⁺ in the cation exchange capacity (CEC) is called "exchangeable sodium percentage" (ESP). At ESP > 15%, the soil is considered solonetzic (White, 2006; Weil, 2017).

3. Clay particle dispersion: Na⁺, having a small charge and large hydration shell, is weakly attracted to the negatively charged clay surface. This leads to expansion of the diffuse layer and repulsion of particles from each other. Clay colloids pass into a sol state (dispersed state), leading to aggregate destruction, compaction, and loss of porosity (Birkeland, 1984; White, 2006).

4. Formation of a dense solonetzic horizon: Dispersed clay migrates downward with percolating water (during periodic rains) and settles in the middle part of the profile, forming a dense, structureless, columnar or prismatic horizon (Bn). It is this dense horizon that makes Solonetz almost impermeable to roots and water (Scheffer et al., 2018).

Diagnostic Features of Solonetz in the Profile

Solonetz have a well-expressed, easily recognizable profile (Foth, 1990; Weil, 2017; Birkeland, 1984):

1. Upper horizon (A or E): Often light, sandy, eluviated, 5–20 cm thick. It is poor in humus and nutrients.

2. Solonetzic horizon (Bn, Bna, or Btna): This is the main diagnostic horizon. It has a characteristic columnar or prismatic structure—vertical peds with rounded "capitals" at the top. The horizon is dense, compacted, clayey, often with signs of gleying (bluish tints). Upon drying, it cracks into large clods. This is where exchangeable Na⁺ content is maximal (ESP often > 30–50%) (White, 2006; Scheffer et al., 2018).

3. Lower transitional horizon (BC or C): Often carbonate, gypsum-bearing, saline, as the lower part of the profile may retain remnants of primary salinization.

4. Reaction: pH of water extract > 8.5, sometimes up to 10.5 and above. This is due to hydrolysis of Na-colloids:

The formed NaOH alkalizes the environment (White, 2006; Weil, 2017).

Morphological Types of Solonetz

Depending on formation conditions, several morphological types of Solonetz are distinguished (Birkeland, 1984; White, 2006; Scheffer et al., 2018):

1. Meadow-steppe Solonetz—form under conditions of periodic wetting (temporary salinization) and subsequent leaching. Characteristic of dry steppe and forest-steppe zones (southern Russia, Ukraine, Kazakhstan). Have a well-developed columnar horizon.

2. Solonetz-Solonchaks—transitional type, where along with exchangeable Na⁺, free salts are also present (usually in the lower part of the profile). With further desalinization, they can transform into typical Solonetz.

3. "Blind" or "crusty" Solonetz—occur in desert and semi-desert regions, where the upper horizon is strongly compacted, often with a surface crust and cracks.

4. Relict Solonetz—ancient Solonetz preserved on surfaces where modern conditions do not favor their formation (e.g., on terraces, in areas with modern humid climates). They represent paleosols.

Conditions Necessary for Solonetzation

For solonetzation to develop, a complex of factors is necessary (Birkeland, 1984; White, 2006; Weil, 2017):

1. Presence of Na⁺ in the soil solution (initial salinization or input from groundwater).

2. Periodic wetting sufficient to leach neutral salts (chlorides, sulfates) from the upper part of the profile but insufficient for complete removal of Na⁺ from exchange positions. This is possible during transition from an arid to a semi-arid climate or during changes in hydrological regime (e.g., after cessation of irrigation).

3. Presence of clay minerals with high cation exchange capacity (predominantly 2:1—montmorillonite, vermiculite, hydromicas), capable of dispersion and swelling.

4. Carbonates in the parent material (CaCO₃, MgCO₃)—they act as a buffer during replacement of Ca²⁺ by Na⁺, as released Ca²⁺ can react with carbonates, maintaining a high level of Na⁺ in solution.

Difference of Solonetzation from Other Processes

Solonetzation is often confused with lessivage and podzolization due to the presence of a light upper and dense lower horizon. However, there are clear differences:

(White, 2006; Birkeland, 1984; Weil, 2017)

Evolutionary Series: Solonchak → Solonetz → Solod

Solonetzation is part of a genetic series of soils reflecting evolution under changing moisture conditions or reclamation (Birkeland, 1984; White, 2006; Scheffer et al., 2018):

1. Solonchak: Accumulation of free salts, structured soils, pH usually < 8.5.

2. Solonetz: Leaching of neutral salts → accumulation of exchangeable Na⁺ → dispersion → formation of columnar horizon, pH > 8.5.

3. Solod (or Solodized Solonetz): With further wetting and leaching, Na⁺ is displaced and removed, acidity increases (pH decreases), the columnar horizon is destroyed, and an acidic, gleyed, sandy profile is formed. This is the final stage of Solonetz degradation.

This series is especially well expressed in soils of southern Russia, Ukraine, Kazakhstan, as well as in Australia (White, 2006; Birkeland, 1984).

Effect of Solonetzation on Plants and Physical Properties

Solonetzation is one of the strongest degradation factors (Weil, 2017; Eash et al., 2016; Foth, 1990):

1. Low water permeability: Due to clay dispersion and swelling of Na-colloids, the soil becomes practically impermeable to water. Water stagnates on the surface without penetrating deeper. This leads to surface waterlogging and development of gleying (see Section 6).

2. Poor aeration: Absence of pores and aggregates leads to oxygen deficiency in the root zone.

3. Toxicity of Na⁺ and OH⁻: High pH and excess Na⁺ in the soil solution cause osmotic stress and specific toxic effects on roots.

4. Nutrition disruption: Na⁺ competes with K⁺ and Ca²⁺ for uptake, causing potassium and calcium deficiency.

5. Formation of a dense crust: Upon drying, the surface of Solonetz is often covered with a hard crust that prevents seed germination.

Reclamation of Solonetz

Restoration of Solonetz is a complex and expensive process. Main methods (Eash et al., 2016; Weil, 2017; Birkeland, 1984):

1. Gypsum application (CaSO₄·2H₂O): This is the main chemical method. Gypsum supplies Ca²⁺, which displaces Na⁺ from the exchange complex. The formed sodium sulfate (Na₂SO₄) is easily leached during washing:

Gypsum doses are calculated based on initial ESP and desired final value (usually 5–10 t/ha or more is required) (Weil, 2017; Eash et al., 2016).

2. Leaching with drainage: After gypsum application, intensive leaching of the soil is necessary to remove the formed neutral salts (Na₂SO₄, NaCl). Good drainage and sufficient water are required.

3. Application of sulfur or sulfuric acid: In some cases (in the presence of carbonates), elemental sulfur (S) can be used, which oxidizes to sulfuric acid, binding carbonates and releasing Ca²⁺, which then displaces Na⁺.

4. Bioreclamation: Cultivation of salt- and sodium-tolerant crops (e.g., some wheat, barley, alfalfa varieties, as well as halophytes) that absorb Na⁺ and contribute to structure improvement through roots.

5. Deep tillage: In some cases, it helps to mechanically break the dense solonetzic horizon, but without chemical reclamation, the effect is short-lived.

Geographic Distribution

Solonetz are widely distributed (Birkeland, 1984; White, 2006; Weil, 2017):

• Dry steppe and semi-desert zones of the temperate belt: Southern Russia, Ukraine, Kazakhstan, Mongolia, northern China, USA (Great Plains, especially North Dakota, Montana), Canada (southern Alberta and Saskatchewan), Argentina (Pampas).

• Tropical and subtropical savanna zones with seasonal wetting: Africa (Sahel, East Africa), India (provinces with saline alluvium), Australia (extensive areas).

• Irrigated agricultural areas with improper reclamation (secondary solonetzation)—Central Asia, China, Egypt, Pakistan, USA (California Valley) (Weil, 2017).

In soil classification systems, solonetzation is a diagnostic feature for the identification of Solonetz (WRB) and Natric subgroups in various orders (Alfisols, Aridisols, Mollisols) in Soil Taxonomy (Buol et al., 2011; IUSS Working Group WRB, 2014).

Definition and Essence of the Process

Carbonatization is the process of accumulation in the soil profile of calcium and magnesium carbonates (primarily CaCO₃, as well as CaMg(CO₃)₂—dolomite), formed as a result of secondary precipitation from the soil solution entering with atmospheric precipitation, groundwater, or released during weathering of carbonate-containing rocks (Birkeland, 1984; Weil, 2017).

Unlike salinization, where readily soluble salts (chlorides, sulfates) accumulate, carbonatization is a process of accumulation of sparingly soluble salts of carbonic acid. Calcium carbonate has a solubility of only about 0.001 g/100 g water (at 25 °C), which is hundreds and thousands of times less than NaCl or Na₂SO₄ (Birkeland, 1984; White, 2006). Therefore, carbonates accumulate in soils already with a small moisture deficit—in semi-arid and subhumid conditions, while accumulation of readily soluble salts requires a drier climate (arid and extra-arid) (Scheffer et al., 2018).

Carbonatization is a process opposite to leaching. In humid (wet) conditions, carbonates are completely leached from the soil. In arid and semi-arid conditions, they accumulate, forming carbonate horizons (Bk, Ck, K) —one of the most common diagnostic features of soils in dry regions (Foth, 1990; Weil, 2017).

Chemical Basis: Carbonate System Equilibrium

Carbonatization is the result of a complex chemical equilibrium involving carbon dioxide, water, and carbonates (Birkeland, 1984; Weil, 2017; White, 2006):

1. CO₂ dissolution in soil air: CO₂ is constantly released in the soil through root and microbial respiration. The partial pressure of CO₂ in soil air can be 10–100 times higher than atmospheric (up to 1–3 kPa vs. 0.04 kPa in the atmosphere). CO₂ dissolves in soil moisture, forming carbonic acid:

2. Carbonate dissolution: Carbonic acid reacts with carbonates (e.g., CaCO₃), converting them to soluble bicarbonates:

This reaction proceeds to the right at high CO₂ content (in the upper, biologically active part of the profile) (White, 2006; Weil, 2017).

3. Transport: Soluble bicarbonates Ca(HCO₃)₂ and Mg(HCO₃)₂ migrate with downward water flow to the lower horizons.

4. Carbonate precipitation: When the partial pressure of CO₂ decreases (below the zone of active root activity), with pH increase, or with water evaporation, the equilibrium shifts to the left, and carbonates precipitate:

It is in this way that the carbonate accumulation horizon (Bk or Ck) is formed (Birkeland, 1984; Scheffer et al., 2018).

Conditions Necessary for Carbonatization

For active carbonatization, a combination of the following factors is necessary (Birkeland, 1984; Weil, 2017; White, 2006):

1. Source of Ca²⁺ and Mg²⁺. These can be:

• Carbonate parent rocks (limestones, dolomites, marls, carbonate moraines, loess).
• Weathering of calcium- and magnesium-containing silicates (plagioclases, amphiboles, pyroxenes) in the upper part of the profile.
• Atmospheric precipitation and dust (especially in arid regions, where dust is rich in carbonates).
• Groundwater saturated with Ca(HCO₃)₂.

2. Semi-arid or subhumid climate with positive or near-zero water balance (evapotranspiration ≥ precipitation). In these conditions, the downward water flow is sufficient for dissolution and transport of bicarbonates but insufficient for their complete leaching beyond the profile.

3. Sufficient biological activity for CO₂ production in the upper part of the profile (for carbonate dissolution).

4. Presence of carbonates in the underlying horizons or in groundwater—to provide "seeding" during precipitation.

Sources of Carbonates in Soils

Carbonates accumulating in soils can have different origins (Birkeland, 1984; Weil, 2017; Scheffer et al., 2018):

1. Lithomorphic (inherited): Carbonates remaining from the parent material (e.g., during incomplete leaching of limestones). They are distributed unevenly through the profile, often in the form of fragments or concretions.

2. Pedogenic (secondary): Carbonates precipitated from the soil solution during carbonatization. It is these that form diagnostic carbonate horizons (Bk, K). They have characteristic morphology—thin films, pseudomycelium, concretions, continuous layers (Birkeland, 1984).

3. Atmospheric-dust: Carbonates entering with wind dust (especially in deserts and semi-deserts, where up to 100–200 kg/ha/year of CaCO₃ is deposited). In some regions (southern USA, North Africa, Central Asia), this source may be dominant (Weil, 2017; Birkeland, 1984).

4. Hydromorphic: Carbonates precipitated from groundwater during its capillary rise and evaporation. This mechanism is especially important in floodplains, deltas, and closed depressions.

Morphological Stages of Carbonatization

Carbonatization is a long-term process that passes through a series of stages, each with characteristic morphological features (Birkeland, 1984; Gile et al., 1966; Weil, 2017):

Stage I (initial): Individual thin coatings (films) of carbonates appear on the underside of stones and pebbles (in coarse-textured soils) or single thin threads (pseudomycelium) in the matrix (in fine-textured soils). CaCO₃ content is usually < 5%.

Stage II (moderate): Carbonate films become continuous, covering most stones; individual small concretions (up to 1–2 cm) appear in the matrix. CaCO₃ content—5–15%. Soil whitening begins to appear.

Stage III (developed): Films on stones become thick (up to several millimeters), often with internal layers; concretions enlarge, their number increases. The matrix is significantly whitened (carbonates occupy 30–70% of volume). CaCO₃ content—15–25%.

Stage IV (intense): A continuous carbonate horizon (K or petrocalcic) forms—carbonates cement the matrix and stones into a dense mass. Lamellar (layered) layers appear—thin, often horizontal, dense carbonate crusts formed as a result of periodic flooding and evaporation. CaCO₃ content—25–50% or more (Birkeland, 1984; Weil, 2017).

Stages V–VI (calcrete): A thick, dense, almost monolithic calcrete plate forms—a carbonate crust up to several meters thick, often with signs of redeposition, brecciation, and secondary cementation. CaCO₃ content can reach 80–90% (Gile et al., 1966; Birkeland, 1984).

Diagnostic Carbonate Horizons

In soil classification systems, two main diagnostic carbonate horizons are distinguished (Birkeland, 1984; Buol et al., 2011; IUSS Working Group WRB, 2014):

1. Calcic horizon (Bk or Ck): A horizon with clear secondary carbonate accumulation but not yet cemented into a monolith. CaCO₃ content > 15%, and at least 5% more than in the underlying horizon. Thickness at least 15 cm.

2. Petrocalcic horizon (K or Bkm): A cemented (hardened) carbonate horizon—calcrete plate. It is so dense that it does not slake in water and cannot be penetrated by roots. Thickness at least 10 cm. It acts as a water-restricting layer, which can cause perched water and gleying above in the profile.

Relationship with Other Processes: Place of Carbonatization in the Accumulation Series

Carbonatization is only one stage of the accumulative series in arid and semi-arid soils. With depth, as moisture decreases and salt concentration increases, increasingly soluble compounds accumulate sequentially (Weil, 2017; Birkeland, 1984; White, 2006):

Upper part of the profile (most leached): Accumulation of relatively immobile compounds—SiO₂ (opal, chalcedony)—podzolization (in acidic conditions) or silicification (in alkaline conditions).

Middle part (carbonatization zone): Accumulation of CaCO₃ (the least soluble of the common salts)—calcic horizon (Bk).

Below (gypsum and salt accumulation zone): With further decrease in moisture, more soluble gypsum (CaSO₄·2H₂O) accumulates—gypsic horizon (By).

Lowest part (or surface in extremely arid conditions): Accumulation of the most soluble salts—chlorides, sulfates of sodium and magnesium—saline horizon (Bz, salic).

This series reflects the sequential increase in solubility: CaCO₃ (0.001 g/100 g) → CaSO₄·2H₂O (0.2 g/100 g) → NaCl (35.7 g/100 g) and Na₂SO₄ (4.8–12 g/100 g) (Weil, 2017; Birkeland, 1984).

Carbonatization and Climatic Zonality

The depth of the carbonate horizon is a reliable climatic indicator (Birkeland, 1984; White, 2006; Weil, 2017):

• In humid regions (precipitation > 600–800 mm/year): Carbonates are completely leached from the profile. The depth to a possible relict carbonate horizon may exceed 2–3 m.

• In semi-arid regions (precipitation 400–600 mm/year): The carbonate horizon lies at a depth of 50–150 cm. These are classical Chernozems and Chestnut soils with a carbonate horizon in the lower part of the profile.

• In arid regions (precipitation 200–400 mm/year): The carbonate horizon rises to the surface—to a depth of 20–50 cm. Soils—brown semi-desert, gray-brown desert.

• In extra-arid regions (precipitation < 200 mm/year): The carbonate horizon may reach the surface, forming calcrete crusts or "desert limestone" (Birkeland, 1984; Weil, 2017).

Ecological and Agronomic Significance of Carbonatization

Carbonatization has a dual effect on soil fertility (Eash et al., 2016; Weil, 2017; Foth, 1990):

Positive aspects:

• Carbonates serve as a source of Ca and Mg for plants (although in poorly available form at pH > 8).
• The carbonate horizon buffers acidity, preventing acidification even during prolonged leaching (therefore, carbonate soils often have a neutral or slightly alkaline reaction).
• At initial stages, carbonatization improves structure (carbonate films cement aggregates), increasing water permeability.

Negative aspects:

• At high carbonate content (especially with calcrete formation), water permeability and aeration deteriorate, and a water-restricting layer forms.
• In alkaline conditions (pH > 7.5–8), availability of micronutrients decreases—Fe, Zn, Mn, Cu, B (they pass into poorly available forms). This leads to typical chlorosis (leaf yellowing due to iron deficiency) and other mineral nutrition disorders in carbonate soils (Weil, 2017; Eash et al., 2016).
• Phosphorus (P) in alkaline conditions is bound into sparingly soluble calcium phosphates (apatites, hydroxylapatites), reducing its availability.

Reclamation of Carbonate Soils

Carbonate soils (especially with calcrete) require special improvement practices (Eash et al., 2016; Weil, 2017; Birkeland, 1984):

1. Deep tillage (gypsum application is ineffective): The carbonate horizon, unlike the solonetzic horizon, does not require chemical reclamation. Mechanical destruction of the dense calcrete plate is sufficient to improve water permeability and root access.

2. Application of acidic fertilizers: To increase the availability of micronutrients and phosphorus in alkaline carbonate soils, physiologically acidic fertilizers (ammonium sulfate, superphosphate) or elemental sulfur are used, which acidify the rhizosphere (Eash et al., 2016; Weil, 2017).

3. Chelation of micronutrients: Application of Fe, Zn, Mn as chelates (organic complexes) allows bypassing the problem of their binding by carbonates.

4. Selection of tolerant crops: Many cereals (wheat, barley, sorghum), alfalfa (due to deep root system), and sunflower grow well on carbonate soils. Potatoes, beans, and many fruit trees (apple, pear) grow poorly due to chlorosis.

Geographic Distribution and Classification

Carbonate soils are widely distributed (Birkeland, 1984; White, 2006; Weil, 2017):

• Dry steppe and forest-steppe zone: Chernozems (southern), Chestnut soils, Serozems—on carbonate loess, marls.
• Semi-desert and desert zone: Brown desert-steppe soils, takyrs, gray-brown desert soils.
• Mediterranean regions: Red and brown soils on limestones (terra rossa, terra fusca).
• Karst landscapes of all climatic zones (where parent material is limestone or dolomite).

In soil classification systems, carbonatization is a diagnostic feature for the identification of Calcisols (WRB), Calcic subgroups in various orders (Aridisols, Mollisols, Inceptisols) in Soil Taxonomy, as well as Calcaric and Calcic qualifiers in WRB (Buol et al., 2011; IUSS Working Group WRB, 2014). The presence of a petrocalcic horizon is the basis for identifying Petrocalcic Calcisols.

Definition and Essence of the Process

Ferralitization (from Latin ferrum—iron and alumen—aluminum) is the process of deep chemical weathering of aluminosilicate minerals under humid tropical and subtropical climates, leading to almost complete destruction of primary silicates, removal of silica (SiO₂) and alkali and alkaline-earth bases (Ca, Mg, K, Na) with relative accumulation in the soil profile of relatively immobile hydroxides and oxides of iron (Fe₂O₃·nH₂O) and aluminum (Al₂O₃·nH₂O), and, in some cases, kaolinite (Birkeland, 1984; White, 2006).

The term "ferralitization" was proposed by French soil scientists to describe the process characteristic of ferralitic soils (Oxisols in Soil Taxonomy, Ferralsols in WRB) —soils that are the final product of long-term (millions of years) weathering under hot and humid climates (Scheffer et al., 2018; Weil, 2017).

Ferralitization is the deepest and most prolonged process of chemical transformation of the mineral part of the soil. As a result of its action, the original rock can be transformed to a depth of 50–100 meters or more, turning into a thick layer consisting almost exclusively of weathering-resistant minerals: iron and aluminum oxides and hydroxides, and (depending on conditions) kaolinite (Foth, 1990; White, 2006).

Chemical Basis: Desilication and Hydrolysis

The key chemical process in ferralitization is desilication—the removal of silica (in the form of soluble silicic acid H₄SiO₄) from the soil profile (Birkeland, 1984; White, 2006; Huang et al., 2012).

Under conditions of abundant moisture and high temperatures, hydrolysis of primary aluminosilicates (feldspars, micas, amphiboles, pyroxenes) proceeds with high intensity:

However, during ferralitization, this process goes even deeper: even the formed kaolinite (1:1-layer silicate) under strong leaching conditions can be destroyed with complete removal of silica:

As a result, gibbsite (Al(OH)₃) —aluminum hydroxide, and goethite (FeOOH) or hematite (Fe₂O₃) —iron oxides and hydroxides, are formed (White, 2006; Weil, 2017; Birkeland, 1984).

Thus, ferralitization includes three sequential stages:

1. Destruction of primary silicates (feldspars, micas, amphiboles) with release and removal of alkalis (K, Na, Ca, Mg) and part of the silica.

2. Formation of secondary clay minerals (kaolinite, halloysite)—at this stage, some silica is still retained in the mineral structure.

3. Destruction of kaolinite and almost complete removal of silica—formation of oxide and hydroxide mineral composition (hematite, goethite, gibbsite).

This process has a directional character: with depth (closer to the parent material), more kaolinite and even primary minerals are preserved; closer to the surface, oxides and hydroxides of Fe and Al predominate (Foth, 1990; White, 2006).

Conditions Necessary for Ferralitization

Ferralitization is a strictly climatically conditioned process. For its active development, a combination of the following factors is necessary (Birkeland, 1984; White, 2006; Weil, 2017):

1. Hot climate with high mean annual temperatures (> 20 °C). High temperature accelerates chemical reactions (Van't Hoff's rule: reaction rates double with a 10 °C temperature increase). This makes the tropics and subtropics most favorable for intensive weathering.

2. Abundant and uniform moisture (precipitation > 1000–1500 mm/year) with a leaching water regime. Constant downward water flow ensures removal of weathering products (silica, alkalis) beyond the profile, preventing their secondary precipitation.

3. Good drainage. Water must freely percolate and be removed; water stagnation (as in gleying) hinders the removal of weathering products and can lead to iron reduction and its removal (rather than accumulation).

4. Long development time. Ferralitization is an exceptionally slow process. Hundreds of thousands to millions of years are required for the formation of a mature ferralitic profile. This is why ferralitic soils are confined to ancient planation surfaces (peneplains) on stable continental shields (Africa, South America, Australia, India) (Birkeland, 1984; Foth, 1990).

5. Acidic or neutral reaction. At high pH (alkaline conditions), silica is less mobile and may precipitate as opal or other forms, hindering complete desilication. Therefore, ferralitization is most intense on acidic and neutral rocks (granites, gneisses, quartzites) but can also develop on basic rocks (basalts) provided long-term leaching (White, 2006).

Morphological Features of Ferralitization

Ferralitization leaves characteristic, easily diagnosable features in the profile (Foth, 1990; Weil, 2017; Birkeland, 1984):

1. Deep, thick profile. The thickness of the ferralitic layer can reach tens of meters. The upper part (the actual soil) is often 1–3 m thick, below is a thick weathering crust (saprolite) of the same thickness.

2. Bright, saturated color. Red, orange-red, yellow, and brown tones predominate, due to the high content of iron oxides and hydroxides (hematite gives red tones, goethite—yellow-brown). Color is often uniform through the profile (weak horizon differentiation).

3. Uniform texture. Characterized by high clay fraction content (often > 50–70%), represented predominantly by kaolinite and Fe and Al oxides. At the same time, the soil often has a pseudosandy structure—small (0.2–2 mm), strong, non-slaking aggregates (concretions) that feel like sand (Weil, 2017; White, 2006).

4. Weak horizon differentiation. Unlike Podzols or lessivated soils, ferralitic soils often have weakly expressed horizontation. The upper humus horizon (A) may be thin; beneath it is a thick, uniform oxide horizon (Bw or Bo), which gradually transitions into saprolite (C). This is the result of a long, unidirectional weathering process, during which the profile becomes "homogenized" (Birkeland, 1984; Buol et al., 2011).

5. Low cation exchange capacity (CEC). Since the main part of clay minerals is represented by kaolinite (1:1) and Fe and Al oxides, which have very low CEC (usually 2–5 cmol(c)/kg), ferralitic soils have extremely low capacity to retain cations (Ca, Mg, K, Na). This makes them very poor and requires a constant supply of nutrients from decomposing organic matter (Foth, 1990; Weil, 2017).

6. Low phosphorus sorption capacity. High content of Fe and Al oxides (especially amorphous) leads to strong phosphorus fixation in the form of sparingly soluble Fe and Al phosphates. This is one of the main agronomic problems of ferralitic soils (Weil, 2017; Eash et al., 2016).

Diagnostic Indicators of Ferralitization

For quantitative diagnosis of the degree of ferralitization, several chemical indices are used (Birkeland, 1984; White, 2006; Weil, 2017):

1. Molar ratio SiO₂/Al₂O₃ (or Si/Al). In fresh rocks, this ratio is usually 3–5 (for feldspars—3, for micas—3–4). During ferralitization, it decreases to 2 (kaolinite) and even to 1–1.5 (with gibbsite formation). The lower the ratio, the stronger the ferralitization. In mature ferralitic soils, this ratio is often < 1.8–2.0 (Birkeland, 1984; White, 2006).

2. Ratio SiO₂/(Al₂O₃ + Fe₂O₃) (Pedro index). Used to assess the degree of desilication: the lower the value, the stronger the removal of silica.

3. Content of "free" Fe and Al oxides (extractable by dithionite or oxalate)—reflects the degree of accumulation of secondary minerals.

Ferralitization and Other Processes

Ferralitization is the final stage of long-term evolution under tropical weathering conditions. It is often superimposed on other processes or precedes them (Birkeland, 1984; White, 2006):

• At initial stages (with less intense weathering), fersiallitization may develop—a process in which 2:1-layer clay minerals (montmorillonite, vermiculite) are preserved and silica is not completely removed. This is characteristic of soils with less intense leaching (subtropics, seasonally wet tropics).

• With increasing moisture, fersiallitization transitions to ferralitization—kaolinite and oxides become dominant. This transition is observed when moving from savannas to humid tropical forests.

• With seasonal waterlogging on ferralitic soils, gleying may develop (reduction of Fe and its removal), leading to the appearance of gray, bluish spots among the red matrix (mottling) (Weil, 2017).

• With long-term erosion, the upper part of the ferralitic profile may be truncated, exposing laterite—a denser, often cemented oxide crust (see Section 11). Thus, ferralitization and lateritization are closely related.

Difference of Ferralitization from Other Processes

Ferralitization is often confused with podzolization (both lead to silica removal) and with lessivage (both produce clayey horizons). However, there are fundamental differences (White, 2006; Birkeland, 1984; Weil, 2017):

Ferralitization and Landscape Evolution

Ferralitization is a process that "ages" landscapes. On ancient planation surfaces (peneplains), not affected by glaciations and tectonic movements, ferralitization can create thick weathering crusts that become an integral part of the geological history of the region (Birkeland, 1984; White, 2006).

Classic examples:

• Brazilian Highlands—thick ferralitic profiles on ancient crystalline rocks.
• Central Africa (Congo Basin, South African Plateau)—ferralitic soils on ancient shields.
• Australia (northwest, Kimberley)—ferralitic weathering crusts, often overlying ancient laterites.
• India (Deccan Plateau)—ferralitic profiles on traps (basalts) (Foth, 1990; Birkeland, 1984).

In these regions, ferralitic soils serve as indicators of landscape stability and long-term (neotectonic) quiescence.

Ecological and Agronomic Significance

Ferralitization leads to the formation of soils that have extremely low natural fertility (Eash et al., 2016; Weil, 2017; Foth, 1990):

Positive aspects:

• Good physical properties—high water permeability, good aeration due to pseudosandy structure.
• High water-holding capacity (due to high clay and oxide content).
• Deep soils with a large root zone volume.

Negative aspects:

• Extremely low CEC—soils almost do not retain cations (Ca, Mg, K). All nutrients must come from organic matter (cycling) or fertilizers.
• Strong phosphorus fixation—up to 80–90% of applied phosphorus is bound into unavailable forms. High doses of phosphorus fertilizers or application as chelates are required.
• Micronutrient deficiency (especially Zn, Cu, Mo) due to their binding by Fe and Al oxides.
• High aluminum content (in acidic ferralitic soils)—aluminum is toxic to many crops (pH < 5.5).
• High susceptibility to erosion—when vegetation cover is destroyed, the fine structure is destroyed, and soils are easily washed away.

For the development of ferralitic soils, the following are required:

• Application of high doses of organic fertilizers (to create a humus reserve that serves as a source of nutrients and improves structure).
• Liming (to reduce Al toxicity and increase P availability).
• Application of micronutrients in chelated form.
• Use of crops capable of efficient P uptake (e.g., some maize varieties, legumes with mycorrhizae).
• Use of conservation tillage systems (zero or minimum tillage to protect against erosion) (Eash et al., 2016; Weil, 2017).

Geographic Distribution and Classification

Ferralitic soils occupy vast areas in the tropical and subtropical belts (White, 2006; Buol et al., 2011; IUSS Working Group WRB, 2014):

• South America: Amazonia, Brazilian Highlands, Guiana Highlands.
• Africa: Congo Basin, West Africa (Guinea, Côte d'Ivoire), East Africa (plateaus not affected by volcanism), South Africa (Limpopo province, Mpumalanga).
• Australia: Northern and eastern regions (Queensland, Northern Territory).
• Asia: India (Deccan), Sri Lanka, Indonesia (Sumatra, Kalimantan), Philippines, New Guinea.

In classification systems, ferralitization is a diagnostic process for the identification of:

• Oxisols (Soil Taxonomy)—soils with an oxic horizon.
• Ferralsols (WRB)—soils with a ferralic horizon.
• Krasnozems, Latosols (in older classifications).

One of the most important diagnostic criteria for ferralitic soils is low CEC of the clay fraction (< 16 cmol(c)/kg clay), which reflects the dominance of kaolinite and oxides over 2:1-layer minerals (Buol et al., 2011; IUSS Working Group WRB, 2014).

Definition and Essence of the Process

Lateritization is the process of formation of dense, often cemented, ferruginous or ferruginous-aluminous crusts and concretions in the upper part of ferralitic profiles under seasonally humid tropical climates (Birkeland, 1984; White, 2006).

The term comes from the Latin later—"brick," as in some regions (especially in India and Southeast Asia) this material has long been used for making building bricks. Laterite crusts were easily cut into blocks in the wet state and after drying in air became as hard as stone (Birkeland, 1984; Weil, 2017).

It is important to emphasize: lateritization is not an independent process but rather a special case and the final stage of ferralitization under conditions of periodic drying of the profile. If ferralitization is the accumulation of Fe and Al oxides in a loose, non-cemented form, then lateritization is their compaction and cementation as a result of periodic drying and (in some cases) subsequent redeposition (White, 2006; Birkeland, 1984).

In modern soil terminology, the term "laterite" is often replaced by more precise concepts: "plinthite" (for soft, non-cemented ferruginous material that can harden upon drying) and "petroplinthite" (for already hardened, cemented material) (Buol et al., 2011; IUSS Working Group WRB, 2014).

Difference of Lateritization from Ferralitization

Since these two processes are closely related, it is important to draw a clear distinction between them (White, 2006; Birkeland, 1984; Weil, 2017):

Mechanism of Lateritization: Three Stages

Lateritization is a multi-stage process involving both chemical and physical factors (Birkeland, 1984; White, 2006; Weil, 2017):

Morphological Forms of Lateritization

Laterite formations can have various morphological forms (White, 2006; Birkeland, 1984; Weil, 2017):

1. Plinthite (soft form): Soft, non-cemented, ferruginous-clayey material that becomes hard upon drying and irreversibly hardens upon rewetting. In the field, it can be cut with a knife, but after drying it becomes as hard as stone. Plinthite is often found as spots, veins, or individual accumulations in the soil matrix. Important: in modern classification, plinthite is a diagnostic material that can harden upon drying but is not yet a hard crust itself (Buol et al., 2011; IUSS Working Group WRB, 2014).

2. Petroplinthite (hard form): This is already hardened, cemented plinthite. It forms continuous layers, armor, or crusts from a few centimeters to several meters thick. Petroplinthite often has a vesicular, pumice-like structure—voids left after leaching of soluble substances are visible inside. On the surface, it may be covered with glaze (a thin shiny layer)—the result of evaporation and redeposition of ferruginous compounds.

3. Pisoliths (concretions): These are small (0.5–5 cm) rounded or oval concretions composed of layered ferruginous material. They are often found in the upper part of ferralitic profiles and can form a "pisolitic horizon", resembling bunches of grapes in appearance. Pisoliths are formed by repeated redeposition of Fe compounds around a central nucleus (Birkeland, 1984).

4. Laterite armor: A continuous, very hard, often thick (up to several meters) crust covering the surface of ancient planation surfaces. It can be uniform (massive) or layered (consisting of alternating dark and light layers reflecting periods of wetting and drying) (White, 2006).

Conditions Favoring Lateritization

For the formation of laterite crusts, a combination of several conditions is necessary (Birkeland, 1984; White, 2006; Weil, 2017):

1. Hot climate with alternating wet and dry seasons. Precipitation must be sufficient for intensive weathering (ferralitization), but the dry season must be long enough for evaporation and capillary rise of salts.

2. Shallow groundwater or water-restricting layer. Capillary rise of saline waters from lower horizons (or from groundwater) is the main mechanism for concentrating Fe and Al in the upper part of the profile. Therefore, laterite crusts are often confined to zones with shallow groundwater or to water-restricting horizons.

3. Basic or neutral composition of parent material. Rocks rich in iron (basalts, gabbros, basic gneisses) are more prone to lateritization, as they supply more Fe. On acidic rocks (granites), lateritization is also possible, but crusts are usually thinner and less developed.

4. Stable tectonic regime. Laterite crusts form only on very old, stable surfaces not affected by erosion or tectonic movements for millions of years.

5. Good drainage (but not excessive). Water stagnation (as in gleying) prevents lateritization, as it leads to reduction and removal of Fe. Conditions of seasonally variable moisture are optimal.

Diagnostic Features of Lateritization in the Profile

Laterite formations have characteristic features (White, 2006; Birkeland, 1984; Weil, 2017):

1. Presence of a dense, hard crust (armor) in the upper part of the profile or on the surface. The crust is often rusty-brown, reddish-brown, or dark red in color.

2. Vesicular (pumice-like) structure. Numerous voids left after leaching of soluble substances (silica, salts) are visible inside the crust. In the air, the voids are often filled with secondary minerals.

3. Presence of pisoliths and other concretions in the loose matrix above or below the crust.

4. Characteristic layering. Crusts often have horizontal or inclined layering, reflecting successive episodes of accumulation and cementation.

5. Sharp boundary with the underlying loose material. Under the crust, there is often a lighter, bleached (siliceous or clayey) horizon depleted in iron.

6. Absence of signs of gleying (bluish tints) in the crust itself, although there may be signs of seasonal waterlogging below it.

Lateritization and Landscape Evolution

Laterite crusts play an important role in the geomorphological evolution of tropical regions (Birkeland, 1984; White, 2006; Weil, 2017):

• Armor against erosion. The laterite crust protects the underlying rocks from erosion, so thick laterite armors are often preserved on plateau tops and hills, while slopes lacking this protection are deeply dissected.

• Relief inversion. In some cases, laterite crusts formed in depressions (where water accumulated), after uplift of the territory, end up on hilltops, while softer rocks around them are eroded. This forms inverted relief (inversion remnants).

• Paleoclimatic indicator. The presence of laterite crusts in regions with a different climate (e.g., in the temperate belt) indicates past (Paleogene, Neogene) periods of humid tropical climate. These are important paleosol indicators (Birkeland, 1984; Weil, 2017).

Lateritization and Other Processes

Lateritization is closely related to ferralitization but can also combine with other processes (White, 2006; Birkeland, 1984):

• With gleying: With seasonal waterlogging of the lower part of the profile on ferralitic soils, gleying develops, and in the upper part—lateritization. As a result, "mottled" (spotted) lateritic profiles are formed, where red, oxidized areas alternate with gray, reduced ones.

• With lessivage: In some ancient profiles, laterite crusts may be overlain by younger sediments on which lessivated soils develop. Then the profile shows two cycles (bisolum): lower—ancient lateritic, upper—modern (Birkeland, 1984).

• With carbonatization: In semi-arid regions, carbonates may accumulate on laterite crusts, additionally cementing them.

Ecological and Agronomic Significance

Laterite crusts (petroplinthites) have a predominantly negative impact on land use (Eash et al., 2016; Weil, 2017; White, 2006):

Negative aspects:

• Unsuitability for agriculture. The hard crust is practically impermeable to roots, water, and air. Such soils are considered unsuitable for agricultural use without the use of heavy machinery to break the crust.

• Difficulty in construction. Laterite crusts create problems in road and foundation construction, as they are very hard but (unlike rock) can be unstable when wetted.

• Formation of a water-restricting layer. Even if the crust is not continuous, it can create a water-restricting layer, causing perched water and gleying above in the profile.

Positive aspects:

• Construction material. In some regions, lateritic material is used for making bricks (it is easily cut in the wet state and becomes very strong after drying). This is a traditional building material in India, Africa, and Southeast Asia.

• Ore mining. In some cases, laterite crusts are enriched in aluminum (bauxites) or iron (iron ores) and serve as mineral deposit objects.

Lateritization in Modern Conditions

It is important to understand that most laterite crusts formed in past geological epochs—mainly in the Paleogene and Neogene (5 to 65 million years ago), when the climate was warmer and wetter than at present (Birkeland, 1984; White, 2006). These crusts are relict paleosols. Modern lateritization in some regions (e.g., in parts of Amazonia, Central Africa, Southeast Asia) still occurs but much more slowly.

Thus, lateritization is not only a modern but also a historical (relict) process that serves as an important indicator for the reconstruction of paleoclimates and paleogeography (Birkeland, 1984; Weil, 2017).

Geographic Distribution

Laterite crusts and plinthitic soils are widely distributed (White, 2006; Birkeland, 1984; Weil, 2017):

• Tropical Africa: Extensive laterite armors on the plateaus of West, Central, and East Africa (e.g., in Guinea, Côte d'Ivoire, Ghana, Cameroon, the Yoruba Plateau in Nigeria).
• South America: Brazilian Highlands, Guiana Highlands (laterite crusts on ancient planation surfaces).
• India: Deccan Plateau, especially on basaltic traps (classic laterites of India).
• Southeast Asia: Thailand, Laos, Vietnam, Cambodia, Indonesia (on ancient surfaces, especially on basic rocks).
• Australia: Northern and eastern regions (Queensland, Northern Territory, Western Australia).
• Outside the tropics: Relict laterite crusts are also found in the temperate zone (e.g., in Europe—in the Massif Central in France, southern England, Germany (Eifel, Westerwald)—as evidence of warm Paleogene periods) (Birkeland, 1984; Scheffer et al., 2018).

Classification of Soils with Signs of Lateritization

In modern classification systems, the terms "laterite" and "lateritization" are used limitedly (Buol et al., 2011; IUSS Working Group WRB, 2014):

• In WRB: Plinthosols are distinguished—soils with a plinthic horizon within 50 cm of the surface. This horizon contains > 15% (by volume) of plinthitic material that can harden upon drying. In the presence of a petroplinthic horizon—already hardened, cemented—the soil belongs to Petric Plinthosols (IUSS Working Group WRB, 2014).

• In Soil Taxonomy: Plinthic subgroups are distinguished in various orders (Ultisols, Oxisols, Alfisols)—for soils with plinthitic inclusions. In the presence of a continuous petroplinthic horizon, soils are classified as Petroplinthic (Buol et al., 2011).

Thus, modern classification moves away from the term "laterite" in favor of more precise and diagnostically strict concepts of "plinthite" (non-hardened material) and "petroplinthite" (hardened).

Definition and Essence of the Process

Turbation (from Latin turbatio—"mixing," "stirring") is the process of mechanical mixing of soil material, leading to disruption of the original horizontation, mixing of substances from different horizons, and, in some cases, the formation of specific textures and structures (Scheffer et al., 2018; Birkeland, 1984).

Unlike most other elementary soil-forming processes that create horizons (differentiate the profile), turbation is a destructive (homogenizing) process. It works against directed differentiation, mixing what has been separated (Birkeland, 1984; Weil, 2017).

It is important to understand: turbation is not one process but a whole group of processes, united by a common feature—mixing. That is why in soil science, one speaks not of a single process but of types of turbation (cryoturbation, bioturbation, argilliturbation, etc.) (Scheffer et al., 2018; Birkeland, 1984).

Fundamental Difference: The "Destructive" Process

Turbation is often contrasted with "constructive" processes (humification, lessivage, podzolization, ferralitization). However, this opposition is not absolute: in some cases, turbation can promote the formation of specific horizons (e.g., cryoturbation creates characteristic textures in Cryosols, and bioturbation—krotovinas) (Scheffer et al., 2018; Birkeland, 1984).

Main Types of Turbation

Depending on the agent causing mixing, several main types of turbation are distinguished (Scheffer et al., 2018; Birkeland, 1984; Weil, 2017):

Morphological Features of Turbation

Turbation leaves characteristic, easily diagnosable features in the profile (White, 2006; Birkeland, 1984; Scheffer et al., 2018):

1. Disruption of horizontation: Boundaries between horizons become unclear, wavy, tongue-shaped, or completely disappear.

2. Presence of inclusions: Inclusions of material from other horizons appear in one horizon—krotovinas, wormholes, humus trickles down cracks, large rock fragments on the surface (during cryoturbation) (Weil, 2017).

3. Specific structures:

• Slickensides—smooth, shiny slip surfaces.
• Gilgai—microhighs and microlows.
• Stone polygons and spots—during cryoturbation.
• Shrinkage cracks—characteristic of argilliturbation (Birkeland, 1984; Weil, 2017).

4. Depth uniformity: In strongly turbated soils (e.g., in Vertisols), the profile may be homogeneous—without clear horizons, with uniform color and texture throughout the depth.

Relationship of Turbation with Other ESFPs

Turbation rarely acts in isolation. It often:

• Masks or destroys traces of other processes (e.g., cryoturbation can destroy horizons created by lessivage).
• Promotes the development of other processes (e.g., bioturbation accelerates humification by mixing organic matter with the mineral part; cracks during argilliturbation provide pathways for water movement and gleying development).
• Determines the appearance of many soils, especially in extreme conditions—in cold regions (cryoturbation is the leading process in Cryosols), in arid regions (argilliturbation is the leading process in Vertisols) (Birkeland, 1984; Scheffer et al., 2018).

Ecological and Agronomic Significance of Turbation

Turbation has a dual effect on soil fertility (Eash et al., 2016; Weil, 2017; Foth, 1990):

Positive aspects:

• Bioturbation improves aeration, drainage, mixing of organic matter with the mineral part, which accelerates humification and increases fertility.
• Argilliturbation (in Vertisols) can create deep cracks through which water and roots penetrate into the lower horizons.
• Cryoturbation (in tundra soils) promotes drainage and mixing to a limited depth.

Negative aspects:

• Turbation can destroy horizons created by slow processes (e.g., lessivage), which reduces differentiation and, possibly, the ability of the soil to retain nutrients.
• Argilliturbation in Vertisols leads to difficulty in tillage (soils are sticky, heavy, with low bearing capacity).
• Cryoturbation in tundra regions makes soils unsuitable for most agricultural crops due to an extremely short growing season and harsh conditions.

Definition and Essence of the Process

Bioturbation (from Greek bios—"life" and Latin turbatio—"mixing") is the process of mechanical mixing and redistribution of soil material under the action of living organisms—soil fauna, microorganisms, and root systems of plants (Birkeland, 1984; White, 2006).

Bioturbation is a special case of turbation but is distinguished as an independent ESFP due to its exceptional role in soil formation. If other types of turbation (cryoturbation, argilliturbation) act under specific conditions (cold, clayey soils), then bioturbation is a universal process occurring in any soil where there is life (and life exists almost everywhere) (Weil, 2017; Scheffer et al., 2018).

Bioturbation is figuratively called the "biological mixer" of the soil. It not only mixes but also:

• Changes structure and porosity.
• Redistributes organic matter.
• Accelerates or slows down other ESFPs.
• Creates unique microsites with special properties.

As White (2006) notes, "bioturbation is one of the most powerful and ubiquitous processes shaping the appearance of the soil profile, especially in the upper horizons."

Main Agents of Bioturbation

Bioturbation is carried out by diverse organisms. Let us divide them into main groups (Birkeland, 1984; Weil, 2017; Scheffer et al., 2018):

Mechanisms of Bioturbation: How Organisms Mix Soil

Bioturbation is carried out through several mechanisms (Birkeland, 1984; White, 2006; Weil, 2017):

1. Passing soil through the digestive tract (worms, insects, some vertebrates). This leads to:

• Change in structure and aggregation.
• Enrich