Soil Colloids

1.1. Definition and size boundaries

The term "colloid" comes from the Greek words kolla—glue, and eidos—form. It was first introduced into science by the English chemist Thomas Graham in 1861 while studying substances that did not pass through semipermeable membranes and did not form crystals. However, soil colloids are not just "non-crystallizing substances," but a special category of particles.

In soil science, colloids are generally considered to be particles ranging in size from 1 to 1000 nanometers (1 nm = 10⁻⁹ m) in at least one dimension (Goldberg et al., 2012). For comparison: the diameter of a human hair is about 80,000 nm, a bacterial cell is 1000–2000 nm. The upper limit of the colloidal range is often aligned with the size of clay particles—2 micrometers (µm), which corresponds to 2000 nm (Weil, 2017; White, 2006).

In soil, colloidal particles include:

• mineral colloids — clay minerals;
• organic colloids — humic substances;
• organo-mineral colloids — complexes between them.

1.2. Why size is critical: specific surface area

Imagine a cube with an edge of 1 cm. Its total surface area is 6 cm². If we divide this cube into cubes with an edge of 1 mm, each small cube will have a surface area of 0.06 cm², and the total surface area of all 1000 cubes will be 60 cm²—10 times larger. Continue grinding to 1 µm (0.001 mm). While maintaining a total mass of 1 g, if the particles have a size of 1 µm, their specific surface area will be about 2.6 m²/g. And for particles of colloidal size, for example, 0.1 µm, the specific surface area increases to 26 m²/g (White, 2006).

For clay minerals, which have a platy rather than spherical shape, the specific surface area values are even higher. For example:

According to White (2006) and Weil (2017)

Thus, specific surface area is the surface area per unit mass of a substance. It is the first reason why colloids dominate the chemical processes in soil. The larger the surface area, the more sites for interaction with water, ions, and molecules.

1.3. Colloidal systems: sols and gels

In soil, colloidal particles can exist in two states:

1. Sol — a state where particles are dispersed (separated) in an aqueous medium. In this state, the colloidal system resembles a liquid, with particles suspended and not settling. This state is typical for moist soil, especially after rain, when colloids swell and disperse (White, 2006).

2. Gel — a state where particles are aggregated (joined) into a spatial network that holds water. In this state, the colloidal system resembles a gel or jelly. This state is typical for soil upon drying or for organic horizons rich in humus.

The transition between these states is reversible: soil can swell upon wetting (gel → sol) and harden upon drying (sol → gel). This ability for reversible transformations is critically important for the water and air regime of the soil (Weil, 2017).

1.4. Colloids vs. true solutions and coarse dispersions

To understand the nature of colloids, it is important to distinguish between three types of systems:

Compiled from: White (2006), Weil (2017)

The Tyndall effect is the scattering of light by colloidal particles, which can be observed by passing a beam of light through a solution. This effect allows us to see the "path" of light in a colloidal solution (e.g., in milk, where fat droplets scatter light). In a true solution, there is no such scattering (White, 2006).

The key difference between colloids and coarse suspensions (e.g., a suspension of sand in water) is that colloidal particles do not settle under gravity over a reasonable time—they are acted upon by Brownian motion forces, which keep the particles suspended (Goldberg et al., 2012).

1.5. Why are colloids the basis of the SAC?

So, let's return to the main question. Colloids determine the chemical properties of soil for three reasons:

1. Enormous specific surface area, which serves as the "arena" for all chemical reactions.

2. Electrical charge on the surface, which attracts and retains ions (this is the basis of the SAC—the ability to retain cations and anions in exchangeable form). We will discuss the origin of the charge in detail in section 5.

3. Reactivity of surface groups, which participate in chemical reactions, including forming complexes with metals (including micronutrients), organic molecules (including xenobiotics), and catalyzing transformations of substances (Weil, 2017).

As the authors of the textbook "The Nature and Properties of Soils" figuratively express it, colloids are "the nature and properties of soil in miniature." It is thanks to colloids that soil becomes not just an accumulation of rock fragments, but a geochemical reactor capable of retaining nutrients for plants, purifying water, and resisting pollution (Weil, 2017).

2.1. General principles of structure

All layered silicates (phyllosilicates) are built from two types of elementary sheets:

• Tetrahedral sheet — silica-oxygen tetrahedra (SiO₄)⁴⁻, connected through shared oxygen atoms in a hexagonal network. The central position is occupied by the Si⁴⁺ ion (sometimes substituted by Al³⁺), surrounded by four oxygen atoms (Weil, 2017; White, 2006).

• Octahedral sheet — aluminum-hydroxyl (or magnesium-hydroxyl) octahedra, where the central cation (Al³⁺, Fe³⁺, or Mg²⁺) is surrounded by six hydroxyl groups (OH⁻) or oxygen atoms. In dioctahedral minerals (e.g., kaolinite, muscovite), two of the three octahedral positions are occupied; in trioctahedral minerals (e.g., talc, brucite)—all three are occupied (Weil, 2017).

The combination of these sheets determines the mineral type: 1:1 (one tetrahedral and one octahedral layer) or 2:1 (an octahedral layer between two tetrahedral ones). It is the layer type and the nature of isomorphous substitutions that determine the key properties of clay minerals—swelling ability, magnitude and origin of charge, and cation exchange capacity (Eash et al., 2016).

Let's consider the four main groups most common in soils.

2.2. Kaolinite (kaolinite group)

Kaolinite is a classic representative of 1:1 minerals (one tetrahedral layer bonded to one octahedral layer through shared oxygen atoms). Its structure does not expand: strong hydrogen bonds (—OH of one surface with —O of the other) act between the layers, so water and ions do not penetrate the interlayer space (Weil, 2017).

Key characteristics:

• Specific surface area — only 5–40 m²/g, exclusively external (Eash et al., 2016; White, 2006).
• Charge — predominantly variable (pH-dependent), as isomorphous substitution is practically absent. Negative charge arises at crystal edges due to deprotonation of Al—OH hydroxyl groups at pH > 5–6 (Sparks, 2003).
• Cation exchange capacity (CEC) — low, 3–15 cmol(c)/kg (Eash et al., 2016). The main contribution comes from edge OH groups, not permanent charge.
• Plasticity and stickiness — weakly expressed, as particles are relatively large (0.1–5 µm) and do not swell (Weil, 2017).

Distribution and environmental significance: Kaolinite dominates in highly weathered soils of the tropics and subtropics (Oxisols, Ultisols), where intense hydrolysis and leaching of silica have led to the destruction of 2:1 minerals (Eash et al., 2016). Soils with predominant kaolinite are easy to till, have good water permeability, but are poor in exchangeable bases and require careful fertility management (Foth, 1990). Due to the positive charge at edges at low pH, kaolinite actively sorbs anions, especially phosphates (White, 2006).

2.3. Smectites (montmorillonite group)

Smectites are 2:1 minerals with a high capacity for isomorphous substitution. The most famous representative is montmorillonite, where Mg²⁺ (sometimes Fe²⁺) replaces Al³⁺ in the octahedral layer. This substitution creates a permanent negative charge that does not depend on pH (Weil, 2017; Eash et al., 2016).

Key characteristics:

• Specific surface area — 600–800 m²/g, with up to 80% being the internal surface of interlayer spaces. Nonpolar molecules (N₂) do not penetrate there, so polar liquids (ethylene glycol) are used to measure the total surface area (White, 2006).
• Swelling — interlayer spaces expand upon wetting, as hydrated cations (Ca²⁺, Na⁺) push the layers apart. This leads to significant changes in soil volume upon wetting-drying cycles (Weil, 2017).
• CEC — high, 80–120 cmol(c)/kg, with most of the charge being permanent and available for exchange (Eash et al., 2016).
• Plasticity and stickiness — very high; smectitic soils (Vertisols) are difficult to till, prone to forming deep cracks upon drying (Foth, 1990).

Distribution: Smectites predominate in soils of arid and semiarid regions (Vertisols, some Mollisols), as well as in temperate zones on carbonate parent materials. They provide high fertility due to their large CEC but require careful water management (Weil, 2017). Due to swelling, smectitic soils are unsuitable for building foundations without special preparation (Eash et al., 2016).

2.4. Illite (hydromica group)

Illite is a 2:1 mineral close in structure to mica (muscovite) but with less potassium and more water. Its feature is the presence of non-exchangeable potassium in the interlayer spaces, which is firmly fixed in the hexagonal cavities of the tetrahedral sheets and does not participate in cation exchange (White, 2006; Eash et al., 2016).

Key characteristics:

• Specific surface area — 100–200 m²/g, mainly external, as the interlayer space is partially blocked by potassium (White, 2006).
• Charge — permanent, arising from isomorphous substitution of Al³⁺ for Si⁴⁺ in the tetrahedral layer (this gives a higher local charge density than in smectites). Part of the charge is neutralized by fixed K⁺, so the accessible CEC is less than the total (Eash et al., 2016).
• CEC — 15–40 cmol(c)/kg, but it can vary depending on the degree of weathering and the accessibility of interlayer positions (White, 2006).
• Swelling — limited, as fixed K⁺ pulls the layers together (Foth, 1990).

Distribution: Illite is widespread in temperate zone soils, especially in glacial deposits, loess, and ancient sedimentary rocks. It is an important source of potassium for plants, although not all potassium is immediately available (Eash et al., 2016). Soils with illite have medium CEC, good structure, and moderate plasticity (Weil, 2017).

2.5. Chlorite

Chlorite is a 2:1:1 mineral: an additional hydroxide layer (usually brucitic Mg(OH)₂ or gibbsitic Al(OH)₃) is located between the 2:1 layers, covalently and hydrogen-bonded to the tetrahedral surfaces (Weil, 2017; White, 2006).

Key characteristics:

• Specific surface area — 25–150 m²/g, external, as the hydroxide layer firmly cross-links the packets (White, 2006).
• Charge — mixed: the negative charge from isomorphous substitution in the tetrahedral layer is partially compensated by the positive charge of the hydroxide layer (due to substitution of Mg²⁺ by Al³⁺). As a result, the net charge is not high (Eash et al., 2016).
• CEC — 10–40 cmol(c)/kg, but part of the potential is blocked by the hydroxide layer (Foth, 1990).
• Swelling — absent, as interlayer bonds are very strong (Weil, 2017).

Distribution: Chlorites are found in soils developed on basic and ultrabasic rocks under conditions of weak weathering (e.g., in cold or arid climates). They are often present in young soils and in moraine deposits (Eash et al., 2016). Chlorites have little agronomic significance, but their presence may indicate a low degree of weathering (Foth, 1990).

2.6. Comparative table of mineral colloids

For clarity, let's summarize the main parameters in a table:

Compiled from data from Weil (2017), White (2006), Eash et al. (2016)

2.7. Summary

Mineral colloids are secondary clay minerals whose structure determines their physico-chemical properties. The two main groups—1:1 (kaolinite) and 2:1 (smectites, illite, chlorite)—differ in swelling ability, specific surface area, and predominant charge type. Kaolinite, with its variable charge and small surface area, dominates in highly weathered soils, while smectites and illite, with their permanent charge and high CEC, are characteristic of less weathered, often more fertile soils in temperate and arid climates.

Knowledge of colloid mineralogy allows predicting soil behavior under irrigation, fertilization, liming, and tillage. In the next section, we will move on to organic colloids—humic substances—which, despite their small mass fraction in the soil, often play a decisive role in its chemical activity.

3.1. Origin and nature of organic colloids

The primary source of soil organic matter is plant photosynthetic production. Senescent tissues—leaf litter, roots, stems—reach the surface or directly enter the soil, where they undergo decomposition by soil microorganisms and fauna. During this process, part of the carbon is mineralized to CO₂, part transforms into dissolved organic matter, and the remaining fraction—most resistant to further decomposition—is converted into humic substances (Baldock, 2012; Weil, 2017).

Humus (in a broad sense) is the totality of all organic compounds in the soil, excluding undecomposed plant and animal residues and living biomass (Sparks, 2003). It is the humic substances that represent the colloidal fraction of organic matter (particle size < 1 µm) and possess all the characteristics of colloids: enormous specific surface area, electrical charge, and swelling ability (Eash et al., 2016).

3.2. Chemical composition and functional groups

The average elemental composition of humic substances (calculated on a dry organic matter basis) is as follows (Sparks, 2003; Weil, 2017):

However, more important than the gross composition is the presence of functional groups that determine chemical activity. Among them are the key ones (Sparks, 2003; Scheffer et al., 2018):

1. Carboxyl (–COOH) — the strongest acidic groups, pK_a ≈ 2–5. They dissociate already at pH > 4, creating a negative charge.

2. Phenolic (–OH aromatic) — weaker acids, pK_a ≈ 9–10, dissociate at pH > 8.

3. Alcoholic (–OH aliphatic) and enolic groups — contribute to charge at high pH.

4. Carbonyl (C=O) — participate in hydrogen bonding and complexation.

5. Amine (–NH₂) and amide (–CONH–) — basic groups capable of protonation at low pH, but their proportion is small compared to acidic ones.

6. Quinoid structures — participate in redox reactions.

It is the carboxyl and phenolic hydroxyl groups that are the main sources of variable (pH-dependent) negative charge in humus (Sparks, 2003). Upon increasing pH, deprotonation of these groups increases the negative charge, leading to an increase in cation exchange capacity (CEC) in an alkaline environment (Weil, 2017).

3.3. Classical fractions and the modern view

Traditionally, humic substances are subdivided by solubility in acids and alkalis into three fractions (Sparks, 2003; Eash et al., 2016):

• Humic acids (HA) — soluble in alkalis, insoluble in acids (precipitate at pH < 2). They have a high molecular weight (from 2000 to 1,000,000 Da), contain many aromatic rings and carboxyl groups.
• Fulvic acids (FA) — soluble in both alkalis and acids. Molecular weight is lower (300–2000 Da), oxygen content is higher, and aromaticity is lower. They are more mobile in the soil profile.
• Humin — the fraction insoluble in alkalis, closely associated with the mineral part. It represents the most humified and stable form.

The modern view on the nature of humus has changed significantly in recent decades. The classical concept of humic substances as "macromolecular polymers" has been criticized (Lehmann & Kleber, 2015). In their review article in Nature, the authors argue that humus is not a separate class of compounds, but a continuum of organic substances, differing in the degree of transformation, sizes, and nature of interaction with the mineral matrix. According to this concept:

1. Soil organic matter exists as supramolecular associations—relatively small molecules (oligomers, metabolites, lignin fragments, polysaccharides) held together by weak interactions (hydrogen bonds, hydrophobic forces, van der Waals forces), rather than covalent bonds (Lehmann & Kleber, 2015).

2. The spectrum of molecules is continuous—from fresh plant polymers (cellulose, lignin) to highly oxidized and degraded products that no longer belong to known classes of compounds.

3. Humus stability is not so much due to its "recalcitrant" (resistant to decomposition) chemical structure, but rather to physical and chemical protection through interaction with minerals and aggregation (Baldock, 2012; Lehmann & Kleber, 2015). This means that humus is not an eternally fixed material—it is dynamic and can be mobilized when conditions change (pH, redox potential, soil tillage).

This concept does not deny the value of traditional fractions (HA, FA, humin) as operational groups, but cautions against attributing strict molecular definiteness to them (Sparks, 2003; Scheffer et al., 2018).

3.4. Colloidal properties of humus

Despite their complexity and heterogeneity, humic substances exhibit pronounced colloidal properties:

• Enormous specific surface area — estimates range from 20 to 800 m²/g depending on the measurement method (Weil, 2017). The reason is the combination of external surface (from particles of colloidal size) and internal surface (in pores and cavities of macromolecular associations). Humus can adsorb up to 20 times its mass in water (Weil, 2017).

• High cation exchange capacity (CEC) — for humus, it is 150–300 cmol(c)/kg at pH 7–8, significantly exceeding even the values for smectites (Eash et al., 2016; Sparks, 2003). Moreover, the CEC of humus is strongly pH-dependent: at low values (pH < 4), carboxyl groups are protonated, and CEC is minimal; upon increasing pH, dissociation increases, and CEC can double (Sparks, 2003). This makes humus a crucial buffering component of the soil.

• Anion exchange capacity (AEC) — unlike permanently charged minerals, humus has practically no positive charges (at pH > 3, it always carries a net negative charge). However, at very low pH, some amine groups can be protonated, providing positive centers, but their contribution is usually small (Sparks, 2003).

• Complexation with metals — humus is a natural chelator. Carboxyl and phenolic groups coordinate metal cations (Fe³⁺, Al³⁺, Cu²⁺, Zn²⁺, Mn²⁺, Pb²⁺, Cd²⁺) forming stable inner-sphere complexes. The stability of complexes increases according to the Irving-Williams series: Mn²⁺ < Fe²⁺ < Co²⁺ < Ni²⁺ < Cu²⁺ > Zn²⁺ (Sparks, 2003). This phenomenon is critically important for:
• maintaining the availability of micronutrients for plants (humus "stores" them but can release them when needed);
• detoxification of heavy metals and radionuclides (binding reduces their mobility and toxicity) (Eash et al., 2016).

• Hydrophilic-hydrophobic balance — humus contains both polar groups (—COOH, —OH) and hydrophobic moieties (aliphatic chains, aromatic rings). This allows it to sorb not only ions but also nonpolar organic molecules—pesticides, polycyclic aromatic hydrocarbons, petroleum products (Baldock, 2012; Weil, 2017). The sorption mechanisms include both hydrogen bonding and hydrophobic interactions (partitioning) (Sparks, 2003).

3.5. Role of organic colloids in soil

In summary, several key functions of organic colloids can be identified (Baldock, 2012; Weil, 2017; Scheffer et al., 2018):

1. Nutrient function — humus is a reservoir of nitrogen, phosphorus, sulfur, and micronutrients. Mineralization of part of the organic matter provides their supply in plant-available forms.

2. Exchange and buffering function — due to its high and pH-dependent CEC, humus actively participates in retaining exchangeable cations (Ca²⁺, Mg²⁺, K⁺, NH₄⁺) and regulating acid-base equilibrium. Soils with high humus content resist acidification and alkalization better.

3. Structure-forming function — humus binds mineral particles into aggregates, improving porosity, water permeability, and aeration (more details in the section on organo-mineral complexes).

4. Ecological (protective) function — sorption of toxic metals and organic pollutants on humus reduces their mobility and bioavailability, serving as a natural barrier against contamination of groundwater and food chains.

5. Energy function — humus (especially its labile fraction) is an energy source for heterotrophic microorganisms, maintaining soil biological activity.

Thus, organic colloids are a highly reactive, dynamic, and multifunctional part of the soil adsorption complex. Their significance is especially great in the upper horizons, where they often contribute more to CEC than clay minerals, especially in acidic or sandy soils (Eash et al., 2016). However, in real soil, organic and mineral colloids almost never exist in isolation—they form complex organo-mineral complexes, which we will discuss in the next section. It is these complexes that give soil its unique set of properties, distinguishing it from any other natural environment.

4.1. What are organo-mineral complexes?

An organo-mineral complex is a spatial and chemical association in which organic molecules (predominantly humic substances) are bound to the surface of mineral particles (clay minerals, Fe/Al oxides, amorphous silicas) through various types of bonds (Sparks, 2003; Weil, 2017). The degree of interaction varies from weak electrostatic attractions to strong covalent bonds (ligand exchange). The resulting structures:

• protect organic matter from rapid mineralization (Baldock, 2012);
• alter the surface properties of minerals (charge, hydrophilicity);
• increase soil aggregate stability and resistance to erosion (Eash et al., 2016);
• create unique sorption sites for cations, anions, and organic pollutants (Scheffer et al., 2018).

In most mineral soils, 50 to 90% of the organic carbon in the upper horizons is part of OMCs (Weil, 2017). Exceptions are sandy soils with low reactive surface area or organic horizons (forest litter, peat), where humus is weakly associated with minerals.

4.2. Mechanisms of organo-mineral complex formation

Depending on the predominant bond type, several mechanisms are distinguished (Sparks, 2003; Weil, 2017; Scheffer et al., 2018):

4.3. Role of organo-mineral complexes in protecting organic matter

One of the key functions of OMCs is the protection of organic matter from microbial degradation. Organic molecules sorbed onto minerals become inaccessible to extracellular enzymes of microorganisms and for direct consumption. This phenomenon is explained by several factors (Baldock, 2012; Lehmann & Kleber, 2015):

1. Spatial inaccessibility — organic molecules strongly bound to mineral surfaces cannot be attacked by enzymes due to steric hindrance.

2. Conformational change — upon sorption, the humus molecule changes its three-dimensional structure, masking enzyme recognition sites.

3. Creation of "microbial refuges" — OMCs form pores and channels that microorganisms cannot penetrate (especially within aggregates).

As a result, the mean residence time of organic carbon in soil increases from years to decades (for unbound humus) to hundreds and even thousands of years (for humus within OMCs) (Baldock, 2012; Lehmann & Kleber, 2015). This is one reason why soils are the largest terrestrial reservoir of organic carbon.

4.4. Influence of organo-mineral complexes on soil structure

OMCs play a crucial role in forming the aggregate structure of soil. There is a hierarchical model (Weil, 2017; Eash et al., 2016):

1. Microaggregates (diameter < 250 µm) are formed by binding individual clay particles together using organic polymers (polysaccharides, glycoproteins) and cationic bridges. These aggregates are stable and form the basis of the structural framework.

2. Macroaggregates (> 250 µm) are formed by joining microaggregates into larger structures with the participation of roots, fungal hyphae, and coarser organic residues.

The destruction of OMCs (e.g., under intensive tillage, salinization, or sharp pH changes) leads to aggregate breakdown, soil compaction, reduced water permeability, and increased erosion (Eash et al., 2016).

4.5. Organo-mineral complexes as centers of reactivity

Synergistic effects are manifested in OMCs:

• Increased CEC — even on minerals with low inherent CEC (kaolinite, oxides), sorption of humus can significantly increase exchange capacity, as humus introduces its acidic groups (Sparks, 2003).
• Shift in point of zero charge — coating oxides with organic acids shifts their PZC to a more acidic range, as the negative charge of humus dominates over the positive centers of oxides. For example, goethite with a PZC around 8 may, after sorbing humus, become negatively charged already at pH > 4 (Sparks, 2003).
• Creation of specific sorption sites — OMCs can bind metal cations simultaneously through oxygen groups of minerals and carboxyl groups of humus, increasing the strength and selectivity of binding (Weil, 2017).
• pH regulation and buffering — OMCs, combining strong and weak acid groups (mineral and organic), provide a wide range of buffering capacity (Scheffer et al., 2018).

4.6. Dynamic nature of OMCs

It is important to emphasize that OMCs are not static formations. They undergo constant transformations:

• Sorption-desorption — depending on pH, ionic strength, and cation concentration, part of the organic matter can transition from bound state to dissolved (DOC) and back.
• Microbial decomposition — even within OMCs, part of the organic matter remains accessible to microorganisms, especially in the presence of low-molecular-weight organic acids exuded by roots (priming effect) (Baldock, 2012).
• Aggregation and disaggregation — wetting-drying and freezing-thawing cycles lead to the restructuring of OMCs, affecting the availability and protection of organic carbon (Weil, 2017).

Thus, organo-mineral complexes are not just a mixture of organic and mineral colloids, but a qualitatively new formation where the properties of the components are enhanced and complement each other. It is OMCs that give soil its unique combination of fertility, resistance to degradation, and self-purification ability. Understanding the mechanisms of formation and destruction of OMCs is crucial for developing sustainable land use and soil conservation strategies.

In the next section, we will address the key question that connects all types of colloids and their complexes into a unified functional system—the origin of electrical charge and its influence on soil sorption properties.

5.1. Permanent (structural) charge

Permanent charge arises in the crystal lattice of certain clay minerals as a result of isomorphous substitution—the replacement of one cation by another with a similar ionic radius but lower valence (Eash et al., 2016; White, 2006). This substitution occurs during mineral formation and does not depend on pH or the composition of the soil solution.

Mechanism of origin:

In the tetrahedral and octahedral layers of clay minerals, some positions that in an ideal crystal are occupied by high-valence cations (Si⁴⁺, Al³⁺) become occupied by cations with a lower charge (Al³⁺ instead of Si⁴⁺, Mg²⁺ or Fe²⁺ instead of Al³⁺). The number of oxygen atoms and their geometric arrangement remain the same, but the total positive charge of the lattice becomes less than the negative charge of the oxygen anions. An excess negative charge arises, which must be compensated by cations from the surrounding solution (Weil, 2017).

Examples of isomorphous substitution:

• In smectites (montmorillonite) — in the octahedral layer, Mg²⁺ (sometimes Fe²⁺) replaces Al³⁺. This creates a negative charge of 0.2 to 0.6 per half-unit cell (Eash et al., 2016).
• In illite and micas — in the tetrahedral layer, Al³⁺ replaces Si⁴⁺, giving a higher charge density (from 0.6 to 1.0 per half-unit cell) (White, 2006).
• In vermiculite — substitution occurs predominantly in the tetrahedral layer, creating a charge of about 0.6–0.9 per half-unit cell (Weil, 2017).

Characteristics of permanent charge:

• Independent of pH — the charge magnitude remains constant over a wide pH range (from pH 2 to 11) (Sparks, 2003).
• Always negative (at least in soil minerals) — isomorphous substitutions increasing valence (e.g., Al³⁺ instead of Mg²⁺) are rare and have little agronomic significance (Weil, 2017).
• Associated with the basal planes of 2:1 minerals, i.e., external surfaces and internal interlayer spaces (in smectites and vermiculites) (White, 2006).
• Compensated by exchangeable cations (Ca²⁺, Mg²⁺, K⁺, Na⁺, NH₄⁺, etc.), which are located in the diffuse layer or in the interlayer spaces (Sparks, 2003).

Carriers of permanent charge in soil:

• Smectites (montmorillonite, beidellite)
• Vermiculites (both dioctahedral and trioctahedral)
• Illite and other hydromicas
• Chlorites (partially)
• Some zeolites (rarely in soils)

In contrast, kaolinite and iron/aluminum oxides have practically no permanent charge, as isomorphous substitution is either absent or extremely minor in them (Eash et al., 2016).

5.2. Variable (pH-dependent) charge

Variable charge arises on surface functional groups capable of accepting or donating protons (H⁺) depending on the pH of the solution. This type of charge is characteristic of all organic colloids (humus), the edge surfaces of clay minerals (especially kaolinite), and oxides and hydroxides of Fe, Al, Mn (Sparks, 2003; Weil, 2017).

Mechanism of origin:

On the surface of minerals and organic macromolecules, there are hydroxyl (—OH), carboxyl (—COOH), and other groups that can act as weak acids or weak bases. Upon pH change, these groups can:

• Dissociate (donate H⁺) → acquire a negative charge:
• :chemical
• e–OH ⇌ ≡Me–O⁻ + H⁺
• :
• :chemical
• COOH ⇌ R–COO⁻ + H⁺
• :
• Protonate (accept H⁺) → acquire a positive charge:
• :chemical
• e–OH + H⁺ ⇌ ≡Me–OH₂⁺
• :
• :chemical
• NH₂ + H⁺ ⇌ R–NH₃⁺
• :

In the first case, the charge becomes negative; in the second, positive. Moreover, the charge magnitude is directly proportional to the degree of dissociation or protonation, which is determined by the pH of the solution and the dissociation constant (pKa) of the given group (Sparks, 2003).

Examples of reactions on different surface types:

Characteristics of variable charge:

• Strongly pH-dependent — upon pH increase, negative charge increases (due to dissociation of acidic groups), positive charge decreases; upon pH decrease, vice versa (Sparks, 2003).
• Can be both negative and positive — depending on the pH relative to the point of zero charge (PZC) of the given surface (Weil, 2017).
• Associated with peripheral (edge) sites of crystals (in minerals) or with functional groups over the entire surface (in humus) (White, 2006).
• Compensated by both cations (at negative charge) and anions (at positive charge). This means that variable charge can contribute to both cation exchange and anion exchange capacity (Sparks, 2003).

Carriers of variable charge in soil:

• Humic substances (carboxyl, phenolic, amine groups)
• Kaolinite (crystal edges)
• Oxides and hydroxides of Fe, Al, Mn (goethite, gibbsite, birnessite)
• Allophanes and imogolites (amorphous aluminosilicates)
• Edge surfaces of 2:1 minerals (to a lesser extent)

5.3. Point of zero charge (PZC)

A crucial characteristic of a surface carrying variable charge is the point of zero charge (PZC) — the pH value at which the net surface charge (sum of positive and negative charges) equals zero (Weil, 2017; Sparks, 2003).

At pH < PZC, the surface has a positive net charge (protonated groups predominate), favoring anion sorption.

At pH > PZC, the surface has a negative net charge (dissociated groups predominate), favoring cation sorption.

PZC values for major soil components:

According to: Sparks (2003), Weil (2017), Eash et al. (2016)

Practical significance of PZC:

• For soils dominated by Fe/Al oxides (e.g., Ferralsols, Oxisols), PZC may lie in the range 7–9, so in acidic soils (pH < 6) these surfaces are positively charged, leading to high anion exchange capacity (especially for phosphates, sulfates) (Eash et al., 2016).
• For humus, PZC is very low (3–4), so in most soils (pH > 4), humus carries a negative charge, making it a powerful cation exchanger (Sparks, 2003).
• For kaolinite, PZC is about 4–5, so in acidic soils (pH < 5), kaolinite edges may be positively charged, explaining its ability to sorb anions (phosphates) in acidic media (White, 2006).

5.4. Ratio of permanent and variable charge in soils

In real soil, both types of charge are present simultaneously, and their contribution depends on the colloid composition and pH.

• In soils dominated by 2:1 minerals (smectites, illite, vermiculite), permanent charge dominates. For example, in Vertisols or Mollisols, up to 80–90% of the total CEC is provided by permanent charge (Weil, 2017). Variable charge (mainly from humus and edges) may constitute 10–20%, especially in the upper horizons.
• In highly weathered soils (Oxisols, Ultisols) dominated by kaolinite and Fe/Al oxides, variable charge prevails. The CEC of such soils is very low (3–10 cmol(c)/kg) and strongly pH-dependent: acidification lowers it further, liming increases it (Eash et al., 2016).
• In soils with high humus content (e.g., Chernozems, Andisols), variable charge of humus can make a decisive contribution to CEC, especially in the upper horizons. For example, with 5% humus and its CEC of 200 cmol(c)/kg, the contribution of organic matter to soil CEC can reach 10–15 cmol(c)/kg (Weil, 2017).
• In sandy soils with low clay and humus content, total CEC is low, and its pH dependence can be strong due to the predominance of variable charge (on humus and oxides) (White, 2006).

Illustration: Figure 5.1 (simplified diagram) shows how the net charge of a soil with predominantly variable charge changes with pH: at pH < PZC, the surface may be positive (anion sorption); at pH > PZC, negative (cation sorption). For soils dominated by permanent charge, the curve is almost horizontal.

5.5. Significance of charge differentiation for practical soil science

Understanding the nature of charge has direct practical implications:

1. Determining CEC and AEC — for soils with variable charge (acidic, highly oxidized), CEC is measured at different pH values to assess the potential and effective exchange capacity (Sparks, 2003).

2. Liming of acidic soils — increasing pH increases the negative charge on humus and oxides, leading to an increase in CEC and improved cation retention (Ca²⁺, Mg²⁺, K⁺). However, this also reduces the positive charge, which may decrease phosphate sorption (Eash et al., 2016).

3. Fertilizer application — anionic fertilizers (NO₃⁻, Cl⁻) are weakly retained in soils with predominantly permanent negative charge (due to repulsion) but can be sorbed in soils with variable positive charge (at pH < PZC) (White, 2006).

4. Ecological risk assessment — the mobility of heavy metals and anionic pollutants (arsenates, chromates) is highly dependent on pH and charge type. In acidic soils with variable charge, they may be retained more strongly (Sparks, 2003).

Thus, the charge of colloidal particles is a fundamental property determining the soil's ability to retain and exchange ions. Permanent charge, associated with isomorphous substitution, provides a stable background of exchange capacity, especially in 2:1 minerals. Variable charge, arising from the ionization of surface functional groups, gives the soil flexibility and sensitivity to pH changes, which is particularly important in soils rich in humus and oxides.

In the next section, we will consider how surface charge translates into surface reactivity—the ability of colloids to participate in specific chemical interactions with ions, molecules, and organic compounds, including the formation of inner- and outer-sphere complexes, ligand exchange, and chelation.

6.1. Surface functional groups

The basis of surface reactivity lies in functional groups—atoms or groups of atoms capable of chemical interactions. In soil colloids, three main types of such groups are distinguished (Sparks, 2003; Eash et al., 2016):

6.2. Protonation and deprotonation

The key process determining the charge and reactivity of the surface is reversible reactions of proton addition or removal (H⁺). They can be written in general form:

Protonation (addition of H⁺):

Deprotonation (removal of H⁺):

where ≡S denotes the colloid surface (mineral or organic).

These reactions are fully reversible and obey the law of mass action. Their direction and completeness are determined by:

• pH of the solution — at low pH (excess H⁺), protonation predominates (positive charge); at high pH (deficiency of H⁺), deprotonation predominates (negative charge) (Sparks, 2003).
• Strength of the acid group — the lower the pKa of the group, the lower the pH at which it dissociates (deprotonates) (Weil, 2017).
• Ionic strength — high salt concentrations can shield the surface charge and influence protonation constants (White, 2006).

Practical significance: It is due to protonation and deprotonation that the colloid surface can "adjust" to the soil pH. In acidic soils, positive centers (and anion exchange) are activated; in alkaline soils, negative centers (and cation exchange) are activated. This makes the soil a buffering system with a wide range of pH regulation (Scheffer et al., 2018).

6.3. Inner- and outer-sphere complexes

The terms "inner-sphere" and "outer-sphere" complexes were introduced into colloid chemistry to describe how exactly an ion or molecule binds to a colloid surface. The distinction is fundamentally important for understanding the strength and reversibility of sorption (Sparks, 2003; Weil, 2017).

6.4. Ligand exchange

Ligand exchange is a specific case of inner-sphere complex formation, where an anion (ligand) replaces a hydroxyl group (OH⁻) or a water molecule from the coordination sphere of a surface metal cation (Sparks, 2003; Weil, 2017).

Schematically for iron oxide (goethite):

Or in general form:

where L⁻ is the anionic ligand (e.g., phosphate, arsenate, sulfate, carboxylate).

Features of ligand exchange:

• Requires the presence of monodentate hydroxyl groups on the surface (≡Me–OH) that can be substituted.
• Accompanied by the release of OH⁻ into the solution, which can locally increase pH (in the microenvironment) (Sparks, 2003).
• Particularly characteristic of Fe, Al, Mn oxides and the edge surfaces of clay minerals (especially kaolinite) (White, 2006).
• Provides strong and often irreversible sorption of phosphates, arsenates, molybdates, borates, and other oxyanions (Eash et al., 2016).

Agronomic significance: Ligand exchange is the main mechanism for fixing phosphorus fertilizers in soil, especially acidic soils rich in Fe/Al oxides. Applied phosphate is rapidly bound in inner-sphere complexes and becomes poorly available to plants ("phosphate fixation") (Eash et al., 2016). Therefore, the efficiency of phosphorus fertilizers drops sharply in oxide-rich soils, requiring special application methods (localized, in granules, with organic acids) (White, 2006).

6.5. Complexation with metals (chelation)

Humic substances are natural chelators. Chelation (from Greek chele — claw) is a process where an organic ligand (e.g., a humus molecule) binds a metal cation through two or more donor groups, forming a closed ring (chelate) (Sparks, 2003; Weil, 2017).

Which groups participate:

• Carboxyl (—COO⁻) — the main donor centers.
• Phenolic (Ar—OH) — particularly important for forming stable five- or six-membered rings.
• Amine (—NH₂) — can participate in chelation at low pH.

Chelate stability follows the well-known Irving-Williams series for transition metal complexes (Sparks, 2003):

Cu²⁺ > Ni²⁺ > Pb²⁺ > Co²⁺ > Zn²⁺ > Mn²⁺ > Fe²⁺

This means that humus preferentially binds copper (Cu²⁺) at the expense of, for example, manganese (Mn²⁺). For some toxic metals (Hg²⁺, Cd²⁺), chelation is also very effective, which has important environmental significance (Eash et al., 2016; Sparks, 2003).

Factors influencing chelation:

• pH — at low pH (acidic medium), carboxyl groups are protonated and bind metals weakly; upon pH increase, chelation intensifies (Sparks, 2003).
• Presence of competing cations — Ca²⁺ and Mg²⁺ (the main cations of the soil solution) bind less strongly to humus, but can compete for sites (Weil, 2017).
• Metal/ligand ratio — when metal is in excess, part of it may be sorbed non-chelately, as outer-sphere complexes or precipitates (Sparks, 2003).

Environmental significance of chelation:

• Plant nutrition with micronutrients — chelates retain Fe, Zn, Cu, Mn in an available but protected from leaching form. When these elements are deficient, plants release chelating substances (phytosiderophores, organic acids) that extract micronutrients from humus complexes (Sparks, 2003).
• Detoxification of heavy metals — humus binds Pb, Cd, Hg, As into stable chelates, reducing their mobility, phytotoxicity, and risk of entering groundwater (Eash et al., 2016).
• Metal transport — in some cases (e.g., with fulvic acids, having low molecular weight and high acidity), chelates can be mobile and promote metal migration in the profile (e.g., during podzolization) (White, 2006).

6.6. Reactivity and pH

To conclude this section, it is important to emphasize that pH is the main regulator of surface reactivity. With pH changes, the following parameters change (Sparks, 2003; Weil, 2017):

Thus, the reactivity of the colloidal surface is a dynamic, pH-dependent characteristic. Understanding these patterns underlies many agronomic and environmental measures (liming, phosphorus fertilization, phytoremediation of polluted soils) (Eash et al., 2016; Scheffer et al., 2018).

Summary of section 6: The surface reactivity of colloids is determined by the presence of functional groups (hydroxyls on minerals, carboxyls and phenols in humus), which can be protonated or deprotonated depending on pH. This allows colloids to form two types of bonds with ions and molecules: weak outer-sphere complexes (reversible ion exchange) and strong inner-sphere complexes (including ligand exchange and chelation). It is due to this ability that soil can simultaneously serve as a reservoir of nutrients, a filter for pollutants, and a regulator of natural water composition. In the final section of the lecture, we will synthesize everything discussed and answer the main question: why exactly do colloids determine soil properties.

7.1. Colloids—the basis of soil fertility

Soil fertility is largely determined by its ability to retain and supply plants with mineral nutrients. This ability is almost entirely provided by colloids (Eash et al., 2016; Weil, 2017).

How it works:

• Cation exchange capacity (CEC) of colloids retains Ca²⁺, Mg²⁺, K⁺, NH₄⁺, and other cations in an exchangeable but plant-available form. Without colloids, these ions would be quickly leached by atmospheric precipitation, especially in humid regions (Sparks, 2003).
• Anion exchange capacity (AEC) (especially on oxides and in acidic soils) retains phosphates, sulfates, borates, and other anionic nutrients, preventing their loss (White, 2006).
• Buffering capacity of colloids (especially humus and smectites) resists sharp pH changes, maintaining the optimal range for root nutrition and microbial activity (Scheffer et al., 2018).
• Complexation retains micronutrients (Fe, Zn, Cu, Mn) in an available form but protects them from irreversible fixation in poorly soluble minerals (Sparks, 2003).

Numbers confirming the role of colloids:

• In soil with 4% humus and 20% clay, the contribution of organic matter to CEC can reach 40–50% (at pH 7) (Weil, 2017).
• In sandy soils, where the mineral part has almost no charge, it is humus (even at 1–2% content) that provides almost all CEC and, consequently, the retention of nutrients (Eash et al., 2016).

Thus, colloids are the soil's battery, charging with ions from minerals, fertilizers, and organic residues, and discharging upon plant demand through root exudates and ion exchange.

7.2. Colloids determine soil physical properties

Many physical properties of soil—water permeability, aeration, water holding capacity, plasticity, stickiness, and resistance to erosion—are also controlled by colloids (Weil, 2017; Eash et al., 2016).

Water regime:

• Swelling of smectites upon wetting creates micro- and macropores, determining the water-air regime. Upon drying, these same minerals shrink, forming deep cracks (in Vertisols), affecting water exchange and root penetrability (White, 2006).
• Humus, with its enormous hydrophilicity (capable of holding up to 20 times its mass in water), increases the field capacity of sandy soils and reduces their tendency to dry out (Weil, 2017).
• Kaolinite and oxides, which do not swell, provide good water permeability even at high clay content (typical of tropical Oxisols) (Eash et al., 2016).

Structure and aggregation:

• Organo-mineral complexes bind individual mineral particles into aggregates, forming macro- and microstructure. Aggregates, in turn, determine pore space, root penetration, water, and air exchange (Weil, 2017; Scheffer et al., 2018).
• Destruction of aggregates (e.g., under intensive tillage or salinization) leads to compaction, crust formation, reduced infiltration, and increased erosion. This is a direct consequence of the disruption of colloidal bonds (Eash et al., 2016).

Mechanical properties:

• Smectites impart high plasticity and stickiness to soils, making them difficult to till, but allow the formation of dense, impermeable layers (e.g., for pond construction or landfills) (Foth, 1990).
• Kaolinite, conversely, makes soils "friable," easy to till, but more prone to water erosion (White, 2006).

7.3. Colloids—regulators of environmental safety

Under anthropogenic pressure, colloids play the role of a natural filter and barrier against pollution (Sparks, 2003; Eash et al., 2016).

Sorption of heavy metals:

• Humus and Fe/Al oxides bind Pb, Cd, Hg, Cu, Zn, As, and other toxic elements into strong inner-sphere complexes and chelates. This reduces their mobility, phytotoxicity, and risk of migration into groundwater (Sparks, 2003).
• For example, humus can bind up to 20–40% by mass of some heavy metals (calculated on organic matter) (Weil, 2017).

Sorption of organic pollutants:

• Hydrophobic moieties of humus sorb pesticides, polycyclic aromatic hydrocarbons (PAHs), polychlorinated biphenyls (PCBs), and petroleum products. This slows their migration and provides time for microbial degradation (Baldock, 2012; Weil, 2017).
• Charged groups of humus and minerals can bind ionic organic compounds (e.g., some herbicides) through ion exchange (Sparks, 2003).

Regulation of natural water composition:

• Colloids participate in shaping the chemical composition of soil and groundwater through ion exchange, sorption-desorption, and precipitation-dissolution. This is particularly important for buffering acid rain and preventing salinization (Scheffer et al., 2018).

7.4. Colloids—the link between litho-, hydro-, and atmosphere

Colloids lie at the interface of three media (solid, liquid, and gaseous) and act as the interface through which all key exchange processes occur (Weil, 2017; Sparks, 2003).

• Atmosphere → soil: Humus and clays sorb atmospheric precipitation, gases (CO₂, NH₃, SO₂), and aerosols, including acid-forming components, neutralizing them through buffer reactions.
• Soil → plants: Colloids supply ions to the soil solution, from which they are absorbed by roots. When necessary, roots excrete organic acids and chelators, "taking" elements from the colloidal complexes.
• Soil → hydrosphere: Colloids retain ions, preventing their leaching into groundwater. However, in some cases (e.g., with fulvic acids), they can transport metals and organic matter downward through the profile (podzolization) (White, 2006).

7.5. Synergy of colloidal components

It is important to understand that colloids act not in isolation, but as a system. The synergy of mineral and organic colloids, their organo-mineral complexes, creates emergent properties that individual components do not possess (Lehmann & Kleber, 2015; Weil, 2017).

• Humus → increases the CEC of clay minerals, especially kaolinite and oxides, which have low inherent CEC (Sparks, 2003).
• Clays → protect humus from microbial decomposition, increasing its "lifetime" from years to millennia (Baldock, 2012).
• Organo-mineral complexes → create unique sorption sites with combined reactivity, unattainable for humus or minerals separately (Lehmann & Kleber, 2015).

It is this synergy that makes soil not just a mixture of minerals and organics, but an integrated system, self-regulating and adaptive.

7.6. Summary scheme: why colloids are the main "conductors" of soil processes

Let's summarize all the arguments in a single scheme:

Conclusion

Here, we have answered the main question: why do the finest particles determine soil properties. Because:

1. They possess enormous specific surface area, where almost all chemical reactions occur.

2. They carry an electrical charge (permanent or variable), allowing them to retain and exchange ions.

3. They are chemically active—capable of protonation, deprotonation, ligand exchange, complexation, chelation.

4. They form organo-mineral complexes, where the properties of the components combine and enhance each other.

5. They provide buffering against pH changes, salinization, and pollution.

6. They manage water and air regimes through aggregation, swelling, and pore space.

7. They are the interface between the atmosphere, plants, and groundwater, through which all flows of matter and energy pass.

Ultimately, colloids are the "soul" of the soil. They transform a pile of mineral particles into a self-regulating, living system capable of nourishing plants, purifying water, sequestering carbon, and resisting degradation.