Soil Formation Factors

1.1. Origins of the Theory: V.V. Dokuchaev and the "Great Law" of Soil Formation

Before the work of the outstanding Russian geologist and soil scientist Vasily Vasilyevich Dokuchaev, the geological view dominated science and practice—soil was seen merely as loose rock (or "altered" rock), whose properties were determined exclusively by the composition of the underlying strata. Views on fertility were fragmented and often reduced to mystical or purely agrochemical explanations.

A radical turning point came in 1883 with the publication of Dokuchaev’s classic work Russian Chernozem. While studying chernozems—Russia's main treasure—Dokuchaev did not limit himself to describing their chemical composition or physical properties. He applied a comprehensive geographical approach, comparing soil properties with the conditions under which they form.

Dokuchaev’s main conclusion was revolutionary: soil is not merely a geological formation but an independent natural-historical body, possessing its own structure (profile), internal properties, and patterns of distribution. He wrote: Soil is the direct result of the combined, very close, century-long interaction between water, air, land... on the one hand, and plant and animal organisms and the age of the country on the other (Mukha et al., 2003, p. 13).

Thus, V.V. Dokuchaev was the first in the history of science to formulate the concept of soil formation factors. He identified five key factors, the combined action of which determines the type, properties, and fertility of soil:

1. Parent (soil-forming) rock — the basis on which and from which soil forms.

2. Climate — the supplier of heat and moisture, determining the rate and direction of all biological and chemical processes.

3. Plant and animal organisms — the active transformers of mineral mass into soil.

4. Relief — the redistributor of heat, moisture, and solid material on the surface.

5. Time — the period during which the soil develops.

Dokuchaev emphasized that these factors are not isolated but act in complex interaction. Their combination produces a "centuries-old" result—that is, millennia are required for the formation of mature soil. He not only listed the factors but also showed how, for example, climate and vegetation in different natural zones of Russia form regular soil sequences—from tundra gley soils to desert serozems. This was the birth of genetic soil science, where the origin and properties of soil are inextricably linked.

1.2. Why Dokuchaev's Concept is the Beginning of a Systems Approach

Dokuchaev's concept solved the main problem—it provided an explanation for the diversity of soils. The answer to the question "why can completely different soils be found nearby?" becomes obvious: because different combinations of soil formation factors can occur nearby.

• Change of parent rock. On a watershed, dense granite may underlie a stony, thin soil, while in a neighboring hollow, a thick layer of ancient alluvium has formed a fertile floodplain soil.
• Change of relief. On a hilltop (eluvial position) and at its foot (accumulative position)—different moisture conditions, different influx of organic matter and weathering products. As a result, different soil types within a single catena.
• Change of time. Soils on an ancient river terrace, tens of thousands of years old, and on a modern floodplain, aged in centuries, will radically differ in the degree of profile development.

Dokuchaev's concept laid the foundation for the comparative-geographical method in soil science. It became possible: to judge the factors of soil formation from soil properties (paleopedology); to predict soil properties from combinations of factors; to classify soils by their genesis (Weil and Brady, 2017, ch. 2).

Key Summary

Soil is not an inert substance but a "product" of the interaction of five groups of "builders": parent rock, climate, organisms, relief, and time. It was V.V. Dokuchaev who gave us this concept, transforming soil science from scattered observations into a fundamental science.

2.1. The Essence of H. Jenny's Approach: Formalizing "clorpt"

In his fundamental work Factors of Soil Formation (1941), H. Jenny made a bold attempt to translate Dokuchaev's intuitive ideas into the language of exact sciences. He proposed viewing soil formation as a system in which soil (S) or any of its individual properties (s) is a function of several independent variables—the soil formation factors.

In its classic form, the equation is written as:

S (or s) = f (cl, o, r, p, t, …)

Where:

• S — soil as an integral natural body, and s — a specific soil property (e.g., humus content, thickness of the humus horizon, pH);
• f — function, i.e., a regular dependence;
• cl (climate) — climate;
• o (organisms) — organisms (biota);
• r (relief) — relief (topography);
• p (parent material) — parent rock (lithological factor);
• t (time) — time;
• … — ellipsis indicating that in some cases other factors may be added, such as atmospheric dust, volcanic emissions, or human influence, if dominant (Birkeland, 1984, p. 163; White, 2006, p. 87).

This equation, known as "clorpt" (from the first letters of the Latin factor designations), became a kind of "periodic table" of soil science. It does not claim that we can instantly compute all soil properties knowing just five numbers. It asserts something else: soil is deterministic, i.e., its properties are a regular consequence of a specific combination of factors. If factors change—soil changes. If factors are the same—soils will be the same (or at least very similar).

2.2. Why This Equation is the Key to Understanding Soil Diversity

Jenny's equation gives us a powerful methodological tool to explain why completely different soils can exist side by side. If we observe differences in soils, we must look for differences in one or more factors on the right side of the equation.

Let's consider this with examples we have already touched upon:

1. Different parent rocks (p): Adjacent areas may have granite and limestone. The equation takes the form: S₁ = f (cl, o, r, p₁, t) and S₂ = f (cl, o, r, p₂, t). Since p₁ ≠ p₂, then S₁ ≠ S₂. This is a lithosequence (White, 2006, p. 88).

2. Different time (t): Soils on an ancient terrace (t₁) and on a modern floodplain (t₂) will differ, even if all other factors are the same: S₁ = f (cl, o, r, p, t₁) and S₂ = f (cl, o, r, p, t₂). This is a chronosequence (Birkeland, 1984, p. 163).

3. Different relief (r): On the same parent rock, in the same climate, but on a hilltop and at its foot—different soils. This is a toposequence (or catena) (White, 2006, p. 91).

4. Different climate (cl): Soils of the taiga and steppe, formed on identical parent rocks, will differ dramatically. This is a climosequence (White, 2006, p. 88).

Thus, Jenny's equation not only provides a neat formula. It dictates a research strategy: to understand soil genesis, one must study all the factors in a given area. To establish the influence of one factor, one must find sites where the other factors are as constant as possible. Jenny called this approach "functional" (Birkeland, 1984, p. 163).

2.3. Practical Value and Limitations of the Equation

It is important to understand that Jenny's equation is not a computational formula in the sense we use, for example, Ohm's law. We cannot plug in numerical values like "temperature = 15 °C", "precipitation = 600 mm" and get the exact soil type as output. Jenny was well aware of this. The main value of the equation is heuristic (cognitive) and organizing (Birkeland, 1984, p. 166; Buol et al., 2011, p. 90).

Significance:

• Systems view: It compels us to consider soil as an integrated system where everything is interconnected.
• Basis for classification: Hierarchical soil classification systems (such as USDA Soil Taxonomy or WRB) are largely built on diagnostic horizons, which are a direct consequence of the action of these factors.
• Prediction: Understanding factor dependencies allows soil scientists to predict which soils will be encountered in poorly studied areas based on analysis of climatic, geological, and topographic maps.

Limitations (why the equation is not solved and likely never will be in a full sense):

• Interdependence of factors: Factors are not always independent. For example, vegetation composition (o) itself strongly depends on climate (cl) and relief (r). They influence each other.
• Quantitative assessment: Some factors, especially "organisms" and "parent rock", are extremely difficult to describe with one or a few numbers (White, 2006, p. 89). What are "organisms"? Species composition, biomass, productivity, root system depth, chemical composition of litter. These are all important but difficult-to-formalize parameters.
• Polygenesis: The equation assumes factors are constant over time. However, many soils are polygenetic—formed in different periods under different conditions (e.g., under forest in the Pleistocene and under steppe in the Holocene). In such cases, how do we account for the change of factors over time in a single function? (Birkeland, 1984, p. 166)
• Nonlinearity and threshold effects: The soil's response to changes in a factor is often nonlinear. For example, an increase of 100 mm of precipitation in an arid zone can cause a radical restructuring of the soil, while in a humid zone, it may have little effect on its properties (modern additions will be discussed later).

2.4. Solutions for Practical Use: "Keeping Factors Constant"

Jenny proposed an ingenious way to circumvent the problem of unsolvability of the general equation. He proposed solving not the general equation but particular functions, keeping all factors except one constant. This gave rise to the well-known "-sequences":

• Climofunction: s = f (cl)o,r,p,t — studying the dependence of a soil property on climate when parent rock, relief, organisms, and time are the same.
• Biofunction: s = f (o)cl,r,p,t — the influence of vegetation.
• Topofunction: s = f (r)cl,o,p,t — the influence of relief.
• Lithofunction: s = f (p)cl,o,r,t — the influence of parent rock.
• Chronofunction: s = f (t)cl,o,r,p — the influence of time (Birkeland, 1984, p. 163).

It was this approach, rather than an attempt to solve the whole equation at once, that yielded the greatest results. Soil scientists worldwide, using Jenny's "functions" method, obtained a wealth of empirical relationships. For instance, the classic work of Jenny and Leonard (1934) showed how, with increasing annual precipitation on loess, humus and clay content regularly increase, pH decreases, and the depth of carbonate occurrence increases (Scheffer et al., 2018, p. 347).

In the end, Jenny's equation is perhaps the most brilliant compromise in soil science. It does not provide a final formula, but it provides a method that still works today, and a rigorous conceptual framework into which all our knowledge of soil genesis fits.

Key Summary of Part Two:

Jenny's equation translates Dokuchaev's concept into the language of a mathematical function: Soil = f(climate, organisms, relief, parent rock, time). Although this equation is not directly solvable, it is a powerful tool for systematizing knowledge, planning research, and making predictions. The diversity of soils is a regular result of the variation of these five main variables. Next, we will examine each of these factors individually, starting with climate.

3.1. Water as the Main Agent: Precipitation and its Effectiveness

Water is the "lifeblood" of soil formation. It participates in all chemical reactions (hydrolysis, dissolution, hydration), serves as a medium for substance transport, and is essential for biota. However, not all precipitation is equally effective for soil formation. What matters most is not the absolute annual millimeters, but effective moisture—the portion of precipitation that percolates through the soil profile, participates in reactions, and leaches weathering products (Weil and Brady, 2017, Ch. 2, Box 2.1; White, 2006, p. 91).

Factors determining precipitation effectiveness:

1. Total amount and distribution pattern. Precipitation of 600 mm evenly distributed throughout the year will cause more leaching than the same 600 mm falling as downpours over 2-3 months (dry season) (Weil and Brady, 2017, Ch. 2, Box 2.1).

2. Temperature and evaporation (evapotranspiration). In hot climates, most moisture is lost to evaporation, leaving little for deep percolation. Therefore, 600 mm of precipitation in a cold climate may create a leaching regime, while in a hot climate, only a non-leaching one (Foth, 1990, Ch. 16, p. 262; Weil and Brady, 2017, Ch. 2, Box 2.1). The water balance—the ratio between atmospheric precipitation and potential evapotranspiration—is used to assess this balance (Buol et al., 2011, p. 108).

3. Relief and rock permeability. On steep slopes, water runs off without penetrating the soil. On dense clay, infiltration is minimal. Consequently, on different relief elements and different parent rocks, "effective precipitation" will differ, even with the same total amount (Weil and Brady, 2017, Ch. 2, Box 2.1).

Depending on the balance between moisture input and its loss through evaporation and runoff, several types of soil water regimes are distinguished (which later form the basis of classifications):

• Leaching regime: Precipitation significantly exceeds evaporation. Water percolates through the entire profile and reaches groundwater, leaching soluble salts and weathering products. Characteristic of humid areas—taiga, humid tropics. Leads to podzolization and lateritization (Scheffer et al., 2018, p. 346).
• Non-leaching regime: Precipitation is approximately equal to evaporation. Moisture wets the soil but does not reach groundwater. During dry periods, capillary rise of solutions is possible. Characteristic of steppes and forest-steppes, promotes humus accumulation and carbonate (calcium) accumulation in the profile (chernozems) (Scheffer et al., 2018, p. 347).
• Effluent regime: Evaporation significantly exceeds precipitation. Moisture not only fails to percolate downward but also rises by capillarity from groundwater to the surface, evaporating and leaving salts behind. Characteristic of deserts and semi-deserts, forms saline soils (solonchaks) (Scheffer et al., 2018, p. 349).

Examples of precipitation influence on soil properties:

A classic example is the transect across loess in the USA. With increasing precipitation from west to east, the depth of carbonate leaching, clay and organic matter content regularly increase, and pH decreases (Foth, 1990, Ch. 16, p. 262; Scheffer et al., 2018, p. 347). Another example: in India, at an average annual temperature of 24 °C, the humus content in the top horizon increases from 0.4% at 35 mm of precipitation to 4.5% at 3200 mm (Scheffer et al., 2018, p. 348).

3.2. Temperature: The Accelerator of All Reactions

If water is the "lifeblood," then temperature is the "pulse" of soil formation. Temperature determines the rate of all chemical, physical, and, most importantly, biological processes. This is based on the simple Van't Hoff rule: a 10 °C increase in temperature increases the rate of chemical reactions by 2-3 times (Buol et al., 2011, p. 106; Weil and Brady, 2017, Ch. 2, p. 77).

How temperature affects different aspects of soil formation:

1. Weathering intensity. In warm and humid climates, minerals decompose much faster than in cold and dry ones. This leads to the formation of thick weathering crusts, deep transformation of parent rock, and formation of secondary minerals (kaolinite, iron and aluminum oxides) in the tropics. In cold regions, physical weathering (frost cracking) dominates (Weil and Brady, 2017, Ch. 2, p. 75-77).

2. Organic matter balance. Temperature affects both phytomass production (faster growth) and the rate of its decomposition by microorganisms. In warm, humid tropics, decomposition is so rapid that, despite enormous biomass, little humus accumulates in the soil. In the cold tundra, conversely, decomposition is severely slowed, and organic matter accumulates as thick peat horizons (Scheffer et al., 2018, p. 346; Foth, 1990, Ch. 16, p. 263).

3. Seasonality and freezing. In temperate latitudes, not only the mean annual temperature but also its seasonal course is important. Winter freezing halts all processes, and spring thawing can cause physical destruction of structure (cryoturbation) and increased erosion. In the tropics with an isothermic regime (small seasonal fluctuations), biological processes run continuously year-round (Buol et al., 2011, p. 106).

3.3. Interaction of Temperature and Moisture: Climatic Zones

The most powerful effect on soil formation is exerted not by temperature or humidity alone, but by their combined action. The combination of heat and moisture determines the type of biome (vegetation), weathering intensity, and the nature of substance migration.

• Humid tropical climate (hot and humid): → deep chemical weathering → formation of ferrallitic soils (Oxisols, Ferralsols) with dominance of kaolinite and iron/aluminum oxides; poor in bases.
• Temperate continental climate (warm summers, cold winters, moderate precipitation): → formation of sod process → humus accumulation → formation of fertile chernozems and chestnut soils.
• Cold humid climate (taiga): → weak chemical weathering, intense leaching → formation of podzolic soils (Spodosols) with an acidic, depleted upper horizon.
• Arid climate (hot and dry): → evaporation dominates → salt accumulation → formation of solonchaks and gypsiferous soils.

It was this pattern, first noticed by Dokuchaev, that led to the creation of the theory of soil zones (Birkeland, 1984, p. 166; White, 2006, p. 92). The change of climatic zones from north to south or from west to east inevitably leads to a change in zonal soil types, explaining their global diversity.

3.4. Why Climate is Not Just an "Annual Average"

It is important to emphasize that soil formation is influenced not so much by average climatic indicators as by extreme events and regime. Soil is a "record" of long-term processes, but rare catastrophic downpours (e.g., in deserts) or multi-year droughts can make a decisive contribution. Therefore, when analyzing climate as a soil formation factor, we speak of the climatic regime—the typical sequence and duration of wetting, drying, heating, and cooling periods (Buol et al., 2011, p. 102; Weil and Brady, 2017, Ch. 2, Box 2.1).

Key Summary of Part Three:

Climate is the "chief conductor" of soil formation. It determines how much energy and moisture enters the system, sets the rate of all processes, and creates the basis for zonal soil division. The two main components—precipitation (its regime and effectiveness) and temperature (its level and seasonality)—in their combination govern:

• The type and intensity of mineral weathering.
• The rate of synthesis and decomposition of organic matter.
• The direction of substance movement (leaching or accumulation).
• The formation of the soil water regime (leaching, non-leaching, effluent).

This is why climate often acts as the most powerful factor in soil differentiation at global and regional levels. Next, we will move on to the next factor—organisms—which, using climatic resources, actively transform the mineral mass into soil.

4.1. Role of Higher Plants: Biological Cycle and Profile Formation

Green plants are the key group of organisms. They create primary organic mass (phytomass) through photosynthesis, involving carbon, nitrogen, ash elements, and energy in the biological cycle. The impact of plants on soil formation is multifaceted.

Input and Transformation of Organic Matter

Plants are the main supplier of organic matter to the soil. It comes in two forms:

• Aboveground litter: leaves, needles, branches, stems. Its amount and chemical composition vary greatly depending on the type of vegetation (forest, steppe, tundra) (Buol et al., 2011, p. 119; Mukha et al., 2003, Table 1).
• Root residues and exudates: dying roots, root exudates. In the case of herbaceous plants (steppes, meadows), the bulk of organic matter enters precisely as roots, permeating the upper soil layer. This leads to uniform impregnation of the upper horizon with humus and the formation of a thick humus horizon (chernozems) (Weil and Brady, 2017, Ch. 2, p. 73; Foth, 1990, Ch. 16, p. 265).

The chemical composition of litter determines the direction of organic matter transformation. Litter from coniferous species is rich in lignin, tannins, and resins but poor in bases (Ca, Mg) and nitrogen. This favors the formation of an acidic, slowly decomposing forest floor (O-horizon) and the development of the podzolic process. Conversely, litter from broad-leaved species and steppe grasses is rich in bases and nitrogen, promoting rapid decomposition, a neutral reaction, and humus accumulation in the form of the humus-accumulative process (chernozem formation) (Foth, 1990, Ch. 16, p. 264; Weil and Brady, 2017, Ch. 2, p. 74).

Biogenic Weathering and Biological Cycling

Plant roots and associated microorganisms (mycorrhizae) release organic acids, carbon dioxide, and chelating compounds into the rhizosphere. This greatly enhances chemical weathering of minerals (hydrolysis, complexation), releasing nutrients (K, Ca, Mg, P, trace elements). Plants actively extract these elements from deeper horizons and, after dying, return them to the upper layers with litter. This phenomenon is called biological accumulation or the "element pump" (Scheffer et al., 2018, p. 350). Thanks to this, the most valuable nutrients are concentrated in the upper part of the profile. Steppe vegetation, with its deep root systems, creates a powerful "humus pump," accumulating nutrients in the upper horizon, which ensures the exceptional fertility of chernozems (Foth, 1990, Ch. 16, p. 264).

Influence on Physical Properties and Microclimate

• Roots mechanically loosen the rock, create macropores, improving aeration and water permeability. After root death, their channels serve as pathways for water movement and new root growth.
• Vegetation cover protects the soil from erosion, softens the impact of raindrops, reduces runoff and evaporation, creating a special microclimate under the forest canopy or grass cover.
• Different types of vegetation significantly affect soil temperature (shading) and humidity (transpiration), thereby modifying the effect of the climatic factor (Buol et al., 2011, p. 107).

4.2. Role of Microorganisms: Biochemical Reactor of the Soil

Microorganisms are the "invisible front" of soil formation. One gram of fertile soil can contain billions of bacteria, actinomycetes, and fungi, as well as numerous protozoa and algae (Weil and Brady, 2017, p. 25). Their significance cannot be overstated:

• Decomposition of organic matter: Microorganisms are the main mineralizers. They decompose complex organic polymers (cellulose, lignin, proteins) into simple mineral compounds (CO₂, H₂O, NH₄⁺, NO₃^-, PO₄³⁻). Without this process, nutrients would remain locked in plant residues forever, and the biological cycle would cease.
• Humification: Part of the decomposition products does not mineralize but undergoes complex polycondensation reactions, forming humic substances—dark-colored, high-molecular-weight, resistant to further decomposition compounds. It is humus that gives soil its most important property—fertility (the ability to retain and gradually release nutrients, improve structure, water-holding capacity) (Weil and Brady, 2017, Ch. 1, p. 38; Mukha et al., 2003, p. 18).
• Nitrogen fixation: Some microorganisms (free-living and symbiotic nitrogen fixers) can reduce atmospheric molecular nitrogen (N₂) to ammonia form (NH₄⁺), available to plants. This is a unique process without which the reserves of fixed nitrogen in ecosystems would be depleted (Scheffer et al., 2018, Kap. 1).
• Redox processes: Microorganisms participate in the transformation of iron, manganese, and sulfur. They can create both reducing (anaerobic) conditions, leading to gleying, and oxidizing conditions, favoring the formation of oxides (Huang et al., 2012, Ch. 30, p. 30-5).

4.3. Role of Animals: Bioturbation and Mixing

Soil animals (zoofauna)—from earthworms and insects to rodents—exert a powerful mechanical effect on the soil mass. This process is called bioturbation.

• Mixing and structure formation: Earthworms, termites, ants, moles, and gophers move huge volumes of soil. They bring material from lower horizons to the surface (e.g., termite mounds) and, conversely, drag organic matter downward. This leads to mixing (homogenization) of the profile, destruction of some horizons and creation of others, leveling differentiation processes (Weil and Brady, 2017, Ch. 2, p. 78; Foth, 1990, Ch. 16, p. 264).
• Drainage and aeration: Animal burrows and channels serve as additional macropores, improving water permeability and gas exchange.
• Aggregate formation: Passing through the intestines of earthworms, mineral particles and organic residues are glued by mucous secretions, forming strong, water-stable aggregates—coprolites, which are the basis of valuable granular or crumb structure (Huang et al., 2012, Ch. 30, p. 30-6).

The role of bioturbation is especially significant in steppes (chernozems), where the activity of burrowing animals and earthworms creates a thick, humified, and well-structured profile (Mukha et al., 2003, p. 26). In the tropics, termites play a similar role, completely reworking the topsoil over millennia.

4.4. Influence of Vegetation on Soil Diversity: Forest vs. Steppe

A classic example of how organisms determine differences between soils is the comparison of soils under forest and under herbaceous vegetation in similar climatic conditions (e.g., at the forest-steppe boundary) (Weil and Brady, 2017, Ch. 2, p. 73; Foth, 1990, Ch. 16, p. 264; Buol et al., 2011, p. 121).

• Under forest: Litter is concentrated on the surface. Decomposition proceeds in an acidic environment with fungal involvement. Coarse, acidic humus (mor) is formed. Decomposition products (organic acids) actively leach bases, iron, and aluminum from the upper part of the profile, forming a light eluvial (podzolic) horizon (E). Below, at some depth, accumulation of leached substances occurs (illuvial horizon—Bh, Bs, or Bt). The profile is clearly differentiated.
• Under steppe: The bulk of organic matter—roots—permeates the entire thickness of the humus horizon. Decomposition proceeds in a near-neutral environment with bacterial involvement. Soft, base-saturated humus (mull) is formed, which impregnates the upper part of the profile, forming a thick, dark-colored humus horizon. Eluvial-illuvial differentiation is absent; the profile shows a gradual transition from the humus horizon to the parent rock.

4.5. Humans as a Soil Formation Factor

Although we treat humans separately (in modern additions and in the chapter on anthropogenic soils), it is important to emphasize that in the last millennia, humans have become a fully fledged sixth factor of soil formation (Amundson and Jenny, 1991; Huang et al., 2012, Ch. 38). Humans affect soil:

• Directly: tillage, drainage, irrigation, fertilization and amelioration, terracing, construction, creation of technogenic substrates (Scheffer et al., 2018, p. 351; Mukha et al., 2003, p. 24).
• Indirectly: changing vegetation cover (deforestation, introduction of crop rotations), altering water regime, atmospheric pollution (acid rain).

As a result, human activity can either accelerate or slow down natural processes, and sometimes lead to the formation of completely new soils (e.g., Agrozems, Urbanozems, Technozems). It is human intervention that explains many differences in soils of adjacent areas that appear identical in terms of the "classical" five factors.

Key Summary of Part Four:

Organisms are the active force of soil formation. They convert solar energy and atmospheric carbon into organic matter, destroy and restructure minerals, mix and structure the soil mass. The role of organisms is unique:

1. Plants — set the "soil profile": determine the nature of humus accumulation, acidity, depth of biogenic weathering, and type of biological cycle.

2. Microorganisms — ensure the closure of the cycle by carrying out mineralization, humification, and nitrogen fixation.

3. Animals — are the "engineers" of the soil, redistributing matter throughout the profile, creating structure, and improving hydrological regime.

4. Humans — in the modern era, are added as a powerful factor capable of radically altering or even completely recreating the soil.

It is the differences in the composition and activity of organisms (especially plant communities and soil fauna) that often explain why fundamentally different soils can form on the same parent rock and in the same climate. Next, we will move on to the relief factor, which redistributes the influence of all previous factors in space.

5.1. Influence of Altitude: Vertical Zonation

With increasing absolute altitude, there is a regular change in climatic conditions (primarily temperature and, as a rule, moisture). This phenomenon is called altitudinal (vertical) zonality (Buol et al., 2011, p. 112; Scheffer et al., 2018, p. 348).

For every 100 m ascent in mountains, air temperature drops on average by 0.6 °C (by the dry adiabatic lapse rate). Simultaneously, precipitation generally increases (up to certain altitudes). Consequently, in mountains, one can observe a change in soil zones analogous to latitudinal ones, but compressed in space to a few kilometers.

Classic example: In the Caucasus Mountains, from the foothills to the summit on the same parent rock and similar geological conditions, the following regularly replace each other:

• At 10–500 m altitude: chestnut soils and chernozems (steppe zone).
• At 500–1500 m: gray forest soils (forest-steppe) → brown forest soils (Phaeozems, Luvisols).
• At 1500–2500 m: sod-podzolic soils and Podzols (mountain taiga).
• At > 2500 m: mountain-tundra soils (Cryosols) and primitive stony soils (Scheffer et al., 2018, p. 348).

Thus, relief indirectly affects climate through altitude, and through climate—the type of vegetation and, consequently, the nature of soil formation. This is one of the most striking answers to the question "why can nearby—on adjacent slopes of the same mountain—have different soils?".

5.2. Slope Aspect: North vs. South

In the middle and high latitudes of the Northern Hemisphere, the orientation of the slope relative to the cardinal directions—aspect—is of great importance (Buol et al., 2011, p. 112; Weil and Brady, 2017, Ch. 2, p. 81).

South-facing slopes (insolated): receive more solar radiation, heat up more, snow melts earlier, soil dries out more often. As a result:

• Vegetation is more xerophytic (drought-resistant), often sparse.
• Less organic matter accumulates.
• Chemical weathering may be more intense due to higher temperatures (with sufficient moisture), but leaching processes are often weakened due to moisture deficit.
• Soils may be more calcareous, less thick.

North-facing slopes (shaded): receive less heat, moisture evaporates less, soil is wetter and cooler.

• Vegetation is more mesophytic (shade-tolerant, moisture-loving), often forested.
• Accumulation of organic matter and its humification are more active.
• Leaching processes intensify, soils may be more acidic and thicker.

Classic study: In Michigan (USA), on south-facing slopes, soils formed with a more reddish hue and higher clay content in the B-horizon than on north-facing slopes (on the same parent rocks and of the same age). Meanwhile, on north-facing slopes, the depth of the humus horizon and total profile thickness were greater (Scheffer et al., 2018, p. 348, citing Cooper, 1960). This is explained by the fact that on warmer south-facing slopes, clay formation was more active, while on wetter north-facing slopes, leaching penetrated deeper.

In the Southern Hemisphere, the situation is mirrored: north-facing slopes are warmer, south-facing slopes are cooler.

5.3. Slope Steepness: Erosion and Accumulation

Slope angle (steepness) is one of the most important parameters determining the ratio of infiltration and surface runoff, as well as the intensity of erosion (Weil and Brady, 2017, Ch. 2, p. 80; Scheffer et al., 2018, p. 349).

• Steep slopes: water runs off quickly, not having time to infiltrate. Consequently, "effective wetting" sharply decreases. In addition, surface runoff removes fine earth, especially the upper, most fertile layer. Soils on steep slopes are typically young, thin, with an undeveloped profile. Erosion processes dominate over soil formation. Here one often finds primitive stony soils (Leptosols) or skeletal variants (Foth, 1990, Ch. 16, p. 266).
• Gentle slopes and watersheds: water has time to infiltrate, leaching and weathering processes proceed fully, mature profiles form (e.g., Luvisols, Chernozems). Erosion is minimal.
• Convex and concave slopes: convex parts (summits, breaks) are sites of intense erosion; concave parts (footslopes) are sites of accumulation of eroded material (colluvium). As a result, at the foot, a thick layer of redeposited soil accumulates, often humified and well-moistened, which can give fertile soils, but with buried horizons (Huang et al., 2012, Ch. 29, p. 29-14).

5.4. Drainage and Landscape Position: Catena

As we mentioned in the climate chapter, relief determines drainage—the soil's ability to remove excess moisture. The position in the relief (eluvial, transeluvial, accumulative) creates regular hydrological sequences called catenas (White, 2006, p. 99; Weil and Brady, 2017, Ch. 2, p. 81; Huang et al., 2012, Ch. 29, p. 29-5).

Example of a catena (Clarion–Nicollet–Webster) on a glacial landscape in Iowa, USA:

• Summit (Clarion): well-drained soil, oxidized, bright brown tones, with deep carbonate leaching.
• Mid-slope (Nicollet): moderately drained soil, somewhat wetter than at the summit, with signs of gleying (gray color, mottles) in the lower part of the profile.
• Footslope and depression (Webster): poorly drained soil (constantly wet), often with a shallow water table, with pronounced gley features—bluish and grayish tones, reduced iron forms (Huang et al., 2012, Ch. 29, p. 29-5; Foth, 1990, Ch. 16, p. 266).

Thus, in the same hollow, on the same parent rock and in the same climate, we see a sequential change of soils: from well-drained to waterlogged. These differences are a direct consequence of topography (relief).

5.5. Interaction of Relief with Other Factors

Relief does not act in isolation. It modifies all other factors:

• With climate: altitude and aspect change temperature and moisture.
• With parent rock: on steep slopes, deep rock horizons are exposed; on gentle slopes, thick layers of colluvium or alluvium accumulate.
• With organisms: steep slopes are often devoid of soil and vegetation, while footslopes accumulate organic matter and create centers of more luxuriant vegetation (oasis effect).
• With time: on stable watersheds, soils can develop for millions of years, while on actively eroding slopes, they are constantly "rejuvenated," never reaching maturity.

Key Summary of Part Five:

Relief is the "architect" of the soil cover, which spatially redistributes the action of climate, water, organisms, and even material. It creates soil diversity within a single territory because:

1. Altitude → changes climate → zonal shifts in mountains.

2. Aspect → redistributes solar heat → differences in heat and moisture supply of neighboring slopes.

3. Steepness → regulates runoff and erosion/accumulation → thin soils on steep slopes and thick soils at the foot.

4. Landscape position → determines drainage and groundwater level → forms regular catenas (sequences of soils from well-drained to waterlogged).

It is thanks to relief that soils at a distance of a few tens of meters can be as different as those thousands of kilometers apart when climatic zones change. This makes relief a critical factor for understanding the mosaic nature of the soil cover. Now we will move on to the next factor—parent rock—which serves as the "building material" for the soil.

6.1. What is Parent Rock?

Parent rock is the loose, unconsolidated rock that serves as the mineral basis for soil formation. It is important to understand that this is not necessarily the rock that lies at depth (bedrock). Often, soil forms on sedimentary deposits (colluvium, alluvium, loess, glacial deposits) that may have been brought from elsewhere. The properties of these deposits can differ greatly from the underlying bedrock (Buol et al., 2011, p. 91; White, 2006, p. 89).

Key idea: Parent rock is not just a "passive substrate." It actively influences all subsequent stages of soil formation, especially in young soils. With age, the influence of parent rock gradually weakens, giving way to the "overprinting" of climatic and biological effects, but it never completely disappears (Foth, 1990, Ch. 16, p. 260; Weil and Brady, 2017, Ch. 2, p. 61).

6.2. How Does Parent Rock Influence Soil Properties?

The influence of parent rock is multifaceted and manifests at several levels.

Mineralogical Composition (what we have "at the input")

Rock minerals are divided into primary (quartz, feldspars, micas, amphiboles) and secondary (clay minerals, carbonates, iron/aluminum oxides), which may have been brought in with the sediment or inherited from a previous weathering cycle (Buol et al., 2011, p. 91).

• Presence of carbonates (CaCO₃): carbonate rocks (limestones, marls, carbonate moraines, loess) contain calcium. This greatly slows acidification processes (buffers acidity), maintains high base saturation, and favors the formation of neutral or slightly alkaline soils. For example, Chernozems form on loess, while Rendzinas form on limestones (Buol et al., 2011, p. 95; Foth, 1990, Ch. 16, p. 260).
• Presence of silicate minerals: granite, gneiss, basalt—all rich in silicates. The rate of their weathering determines how quickly and which nutrients will be released. Quartz is very stable and acts as an "inert filler." Feldspars (potassium, sodium, calcium) are an important source of bases. Micas (biotite, muscovite) are a source of potassium and iron.
• Iron and magnesium content: rocks rich in ferromagnesian minerals (basalt, gabbro) produce many iron oxides upon weathering, and consequently, the soils often have red or brown hues and are characterized by higher fertility (Buol et al., 2011, p. 92; White, 2006, p. 88).

Granulometric Composition (texture)

The size of the particles making up the rock (sand, loamy sand, loam, clay) determines:

• Water permeability and water-holding capacity: sandy rocks quickly transmit water but hold it poorly; clayey rocks, on the contrary, absorb slowly but store it well.
• Absorption capacity (ability to retain nutrients): clay particles have a large specific surface area, so such soils are more fertile (if not waterlogged).
• Structure formation: loams and clays enriched with organic matter form aggregates; sands are mostly loose (Foth, 1990, Ch. 16, p. 259; White, 2006, p. 91).

Chemical Composition and Nutrient Regime

Rock rich in easily weatherable minerals (e.g., basic igneous rocks) supplies plants with more potassium, calcium, magnesium, phosphorus, and trace elements. Quartz sands poor in feldspars (oxygen-deficient) give low-fertility, base-poor soils, even under favorable climate (Buol et al., 2011, p. 93; Weil and Brady, 2017, Ch. 2, p. 65).

6.3. Weathering Stability of Minerals (Goldich's Series)

Not all minerals decompose at the same rate. The stronger the crystal lattice, the longer the mineral persists in the soil. This is illustrated by the famous weathering stability series (Goldich's series), which (in reverse order, from least to most stable) looks like this:

• Least stable: olivine, calcium plagioclase (anorthite), augite, hornblende, sodium plagioclase (albite).
• Intermediate: biotite, potassium feldspar (orthoclase, microcline), muscovite.
• Most stable: quartz (White, 2006, p. 88-89; Weil and Brady, 2017, Ch. 2, Table 2.2).

What does this mean in practice?

In the early stages of weathering (in young soils), we find a mixture of all these minerals. But over time (in old, mature soils), all the "weak" minerals will have decomposed, leaving only the most stable ones—quartz, iron and aluminum oxides, and some secondary clays (kaolinite). This is why sandy soils in old tropical landscapes consist almost exclusively of quartz (Weil and Brady, 2017, Ch. 2, p. 65; Buol et al., 2011, p. 92).

This series also helps explain why soils formed on granites (rich in stable quartz and feldspars) weather more slowly and often give acidic, low-fertility soils, in contrast to soils on basalts (rich in unstable ferromagnesian minerals), which weather rapidly and give fertile clay soils.

6.4. Why is the Influence of Parent Rock Particularly Noticeable in Young Soils?

In young soils (Holocene, last 10–12 thousand years), the mineralogical and chemical composition of the soil is almost entirely inherited from the parent rock. Soil properties directly depend on what it formed on: loess, moraine, alluvium, or eluvium of bedrock. This makes parent rock the dominant factor in the early stages of soil formation (Foth, 1990, Ch. 16, p. 260; Birkeland, 1984, p. 166).

In old soils (Pleistocene, hundreds of thousands of years and more), the action of other factors (climate, organisms) "overwhelms" the initial signal of the parent rock. Leaching of bases occurs, accumulation of stable secondary minerals takes place, and the rock can be heavily transformed. However, some "inherited" traits, such as overall texture (if extreme, e.g., pure sand or pure clay) or the presence of resistant minerals (quartz), can persist indefinitely.

Example: Soils on ancient terraces of the Amazon, formed on quartz sands, remain extremely poor even after millions of years, unlike neighboring soils on clay shales (Buol et al., 2011, p. 93). The influence of parent rock never completely disappears—it remains the "skeleton" and the "constraining factor" for all other processes.

6.5. Main Types of Parent Rocks (without excessive detail)

Main page: Mineral basis of soils

For general understanding, it is useful to know a few common groups:

• Igneous rocks (granite, basalt) → give soils rich in primary minerals, but with different weathering rates.
• Metamorphic rocks (gneiss, schist) → soils on them are often similar to igneous ones, but with structural peculiarities.
• Sedimentary rocks (limestone, sandstone, shale) → the most widespread on the Earth's surface (about 75% of land). Soils on them strongly depend on impurities (clay particles in limestone, cement in sandstone).
• Quaternary deposits (loess, moraine, alluvium, colluvium, aeolian sands) — the main parent rocks in temperate latitudes. They create the "loose" thicknesses in which intensive soil formation takes place. Loess is a particularly valuable material: it is calcareous, loamy (good combination of sand, silt, and clay), permeable, and fertile (Mukha et al., 2003, p. 15; Buol et al., 2011, p. 98).

Key Summary of Part Six:

Parent rock is the foundation that sets the "starting conditions" for all soil formation. It determines:

1. Mineralogical composition: the set of primary and secondary minerals, their stability, and ability to release nutrients.

2. Granulometric composition: texture (sand, loam, clay), which controls water, air, and thermal regimes.

3. Chemical activity: presence of carbonates, iron, exchangeable bases—all influencing acidity, buffering, and potential fertility.

4. Age-related "fading": the influence of parent rock is maximal in young soils and gradually weakens in old ones, but never completely disappears.

This is why, when analyzing any soil profile, we always start by identifying the parent rock. And this is why two neighboring soils formed on different parent rocks (e.g., granite and loess), even in the same climate and relief, will differ dramatically. Now, after analyzing all five "classical" factors, we move on to the last, most "mysterious" factor—time.

7.1. Absolute and Relative Age of Soil

In soil science, two concepts of age are distinguished, and it is important not to confuse them (Mukha et al., 2003, p. 23; Birkeland, 1984, p. 166).

• Absolute (calendar) age — the time elapsed since the beginning of the formation of a given soil. Measured in years (or millennia, millions of years) from the "zero point"—the moment when the surface was exposed to soil formation (e.g., after glacier retreat, volcanic eruption, alluvium deposition). Absolute age is established by geochronological methods: radiocarbon dating (¹⁴C) of organic matter, optically stimulated luminescence (OSL) for mineral grains, potassium-argon or uranium-thorium dating (Buol et al., 2011, p. 136; Foth, 1990, Ch. 16, p. 257).
• Relative age — the degree of expression of soil properties and the degree of profile development compared to other soils. A young soil in absolute terms may be highly developed if it forms under favorable conditions (e.g., on volcanic ash in the tropics). Conversely, an old soil by calendar age in a dry, cold desert may be very poorly developed (Scheffer et al., 2018, p. 352).

Thus, relative age is an integral indicator of how far soil formation has progressed, while absolute age is simply a calendar date. For the soil scientist in practice, relative age is often more important because it is directly related to the morphology and properties of the soil.

7.2. Rate of Soil Formation: Why Do Some Processes Go Fast and Others Slow?

The rate at which soil develops depends on the intensity of all other factors (climate, biota, parent rock, relief). Under favorable conditions (hot, humid, loose and base-rich parent rock, gentle relief), soil formation proceeds quickly; under unfavorable conditions (cold, dry, dense quartz rock, steep slope)—extremely slowly.

Different soil properties and horizons form at different rates, and this is a key idea for understanding soil evolution (Foth, 1990, Ch. 16, p. 255; Weil and Brady, 2017, Ch. 2, p. 82).

Fast (years – decades):

• Appearance and accumulation of organic matter in the upper horizon, darkening of the A-horizon (humus accumulation). On fresh alluvial deposits or volcanic ash, a noticeable A-horizon can form in 10–20 years (Weil and Brady, 2017, Ch. 2, p. 82; Scheffer et al., 2018, p. 352).
• Formation of a weakly developed B-horizon (Bw) with color and structure change—from 100 to 1000 years.

Medium (hundreds – thousands of years):

• Accumulation of carbonates (CaCO₃) in the lower part of the profile under arid conditions (calcic horizons, Bk) — from a few hundred to several thousand years.

Slow (thousands – tens of thousands of years):

• Formation of a textural (clay) B-horizon (Bt) with noticeable illuvial clay accumulation. The formation of a noticeable Bt requires a minimum of several thousand years, and under temperate climate—10–15 thousand years (Foth, 1990, Ch. 16, p. 255; Birkeland, 1984, p. 168).
• Significant change in mineralogical composition, destruction of primary minerals, and formation of stable secondary ones (kaolinite, iron oxides)—hundreds of thousands of years.

Very slow (millions of years):

• Formation of deep weathering crusts (>10–20 m thick) and ferrallitic profiles (Oxisols) in the tropics.
• Complete leaching of all bases and most of the silica.

Example from chronosequences: On sandy lake terraces of different ages in Michigan (USA), it was shown:

• 2250 years: soil with A and C horizons (Deer Park).
• 3000 years: Bs-horizon appeared (accumulation of sesquioxides) and E-horizon (Rubicon).
• 8000 years: a thick Bhs-horizon with humus and oxide accumulation formed (Kalkaska) (Foth, 1990, Ch. 16, p. 254; Weil and Brady, 2017, Ch. 2, p. 82).

This example clearly demonstrates how, over time, new horizons appear sequentially, and the profile becomes increasingly complex.

7.3. Interaction of Time with Other Factors

Time never acts in isolation. It always interacts with other factors.

• Climate: In a warm, humid climate, the same 10,000 years will lead to much deeper weathering and profile differentiation than in a cold or dry climate. For example, in the tropics, 10,000 years may suffice to form a ferrallitic soil with a thick lateritic horizon, while in the tundra—only a primitive Cryosol (Buol et al., 2011, p. 138).
• Parent rock: On loose, base-rich rocks (loess, basalt), soil develops faster than on dense, poor ones (quartzite, granite). This means that over the same time, soils of different relative ages will form on different rocks (Birkeland, 1984, p. 166).
• Relief: On steep slopes, erosion processes constantly "renew" the surface, washing away weathering products. Therefore, soils on slopes may remain young in relative age for many millennia, while on neighboring watersheds, where erosion is minimal, they reach maturity (Weil and Brady, 2017, Ch. 2, p. 82).

7.4. Old and Young Soils: What are the Differences?

What do young and old soils look like, and by what signs can they be distinguished in the field?

Young soils (e.g., on river floodplains, moraines of the last glaciation, fresh volcanic ejecta):

• Weakly differentiated profile: most often A-C or A-Bw-C (Bw — weakly developed horizon).
• Absence of a distinct illuvial horizon (Bt, Bhs).
• Colors close to the color of the parent rock (adjusted for humus accumulation in the A-horizon).
• Dominance of primary minerals, few secondary clays.
• High content of bases (calcium, magnesium, potassium), if the parent rock contained them.
• High pH (often neutral or slightly alkaline), if the parent rock is calcareous (Foth, 1990, Ch. 16, p. 260).

Old soils (e.g., on ancient terraces, planation surfaces, stable tropical landscapes):

• Thick, clearly differentiated profile with pronounced eluvial and illuvial horizons (E-Bt, E-Bhs, or thick Bw with deep weathering).
• Altered mineralogical composition: dominance of stable secondary minerals (kaolinite, Fe and Al oxides), destruction of primary ones (feldspars, micas).
• Deep leaching of bases: soil acidic, unsaturated, with low cation exchange capacity.
• Often strong coloration (red, yellow hues from iron oxides).
• Presence of thick accumulation horizons (e.g., carbonate horizons in arid regions, or spodic horizons in humid ones) (Birkeland, 1984, p. 168; Buol et al., 2011, p. 137).

7.5. Why is Time Not Just a "Number of Years"?

In soil science, it is important to understand that relative age is not always directly proportional to absolute age. This is because soil formation is not a linear but an asymptotic process. Initially, changes are rapid, then their rate slows down, and the soil approaches a "steady state" where the rate of change becomes negligible (Birkeland, 1984, p. 168; White, 2006, p. 100).

This law of "diminishing returns" holds for many properties: accumulation of organic matter, clay, carbonates. Initially, growth is exponential, then the curve flattens. For example, humus content in Chernozems reaches a plateau within a few thousand years, while clay accumulation in the B-horizon may continue for tens of thousands of years, but increasingly slowly.

Two important implications:

1. Nonlinearity: Two soils differing by a factor of 2 in absolute age (e.g., 5000 and 10000 years) may differ far less in degree of development than two soils differing by a factor of 2 in the initial period (e.g., 100 and 200 years).

2. Threshold effects: At a certain point in time (threshold), a sharp change in the dominant process may occur (e.g., carbonate accumulation gives way to their leaching under climate change, or destruction of primary clays triggers oxide formation). This makes the time series complex for simple extrapolation.

Key Summary of Part Seven:

Time is the "integrator" of all soil-forming processes. It allows:

1. Accumulation of organic matter and formation of a humus horizon (years – centuries).

2. Creation of textural differentiation (eluvial-illuvial horizons) (thousands of years).

3. Radical change in the mineralogical composition of the rock (hundreds of thousands – millions of years).

4. Bringing the soil to a state close to equilibrium (steady state).

Different properties form at different rates, so soils of the same age can have different degrees of development under different conditions. This is why, to explain the diversity of soils encountered in a single landscape, we must always consider not only what factors are acting but also how long they have been acting. Ancient surfaces (e.g., terraces or watersheds) bear the imprint of millennia or even millions of years of evolution, while young surfaces reflect only recent history.

In the next, concluding part of the lecture, we will discuss modern additions to the factorial model—concepts of catenas, threshold changes, feedbacks, and nonlinearity of soil processes—which make this model even more powerful and explanatory.

8.1. Catenas and Toposequences: Spatial Connection of Soils

The concept of the catena, introduced by J. Milne in 1935 in East Africa, was the first important addition to the factorial model. A catena is a regular sequence of soils replacing each other down a slope from the watershed to the foot, formed on a uniform parent rock (or a series of rocks) and under identical climatic conditions, but differing due to relief, drainage, and redistribution processes (Huang et al., 2012, Ch. 29, p. 29-4; Weil and Brady, 2017, Ch. 2, p. 81; White, 2006, p. 99).

Key idea of the catena: soils in the landscape are not isolated but connected to each other by flows of water, dissolved substances, and solid material.

• Eluvial positions (summits, upper slopes): Export of material prevails (leaching, erosion). Soils here are often more acidic, depleted, with contrasting horizons.
• Transit positions (middle slopes): Material moves without accumulation.
• Accumulative positions (footslopes, depressions, floodplains): Accumulation of material brought from above (colluvium, alluvium, salts) occurs. Soils here are often thicker, enriched in organic matter or salts, with a less differentiated profile.

Example of a catena (after Milne, for East Africa):

• Summit: well-drained acidic soil (Acrisol).
• Slopes: soils with signs of erosion and washout.
• Footslope: accumulation of eroded material, more fertile and moist soils (Luvisol).
• Valley floor: waterlogged soil with gley features (Gleysol) (Huang et al., 2012, Ch. 29, p. 29-4).

A catena is not just a "toposequence" in Jenny's terms (where only relief changes). It is a functionally connected system where substances leached from one soil enter another. This explains why soils on a slope cannot be considered in isolation—they are parts of a single landscape-geochemical conjugation.

Modern development of the idea:

The catena concept was extended to the "soil-landscape" model and the watershed (Huggett, 1975). In this model, the basic unit of analysis is not an individual pedon but an entire first-order catchment—from the watershed to the river channel—which represents a naturally closed system for material transport (Huang et al., 2012, Ch. 29, p. 29-6). This allows quantitative modeling of carbon fluxes, salts, and weathering products at the landscape scale.

8.2. Nonlinearity of Soil Formation: Why Changes Are Not Always Smooth

The classical model assumed more or less smooth, monotonic changes in soil properties with changes in factors. However, reality is far more complex: many soil processes exhibit nonlinear behavior, including threshold effects, feedbacks, and hysteresis (Birkeland, 1984, p. 166; Huang et al., 2012, Ch. 30, p. 30-24).

Threshold effects

A threshold is a critical point in factor space or time, upon crossing which the nature of the process or the direction of soil development changes abruptly (Huang et al., 2012, Ch. 30, p. 30-24; Weil and Brady, 2017, Ch. 2, Box 2.1).

Examples:

• Carbonate threshold (textural): In soils on carbonate rocks, as long as calcite is present in the profile, pH stays around 7–8, and leaching of silica and clay formation processes are strongly retarded. Once all carbonates are leached (threshold reached), pH drops sharply, and intensive acidic weathering of silicates begins, forming kaolinite and iron oxides (Buol et al., 2011, p. 93).
• Threshold of organic matter accumulation: In some soils, at a certain level of organic matter input, a sharp transition may occur from a state where decomposition dominates accumulation to a state where accumulation begins to prevail (e.g., when shifting from aerobic to anaerobic conditions).
• Hydromorphic thresholds: Upon reaching a certain duration or level of waterlogging, the redox regime changes abruptly, leading to massive reduction of iron and manganese, loss of their mobility, and formation of gleyic horizons (Huang et al., 2012, Ch. 30, p. 30-21).

Thresholds make soil prediction complex: a small change in a factor (e.g., +50 mm of precipitation) may cause no change until a threshold is crossed, after which the soil may drastically restructure.

Positive and negative feedbacks

Feedback is a process in which a change in some soil property influences the intensity of the processes that created that property. This makes the soil a self-regulating system (Huang et al., 2012, Ch. 30, p. 30-24; Weil and Brady, 2017, Ch. 2, p. 76).

Positive feedback (accelerates the process):

• Example (clay accumulation): Once illuvial clay begins to accumulate in the lower horizon, that horizon becomes less permeable. Lower permeability leads to more frequent water stagnation above it. Water stagnation enhances weathering and clay formation in the overlying horizons, ultimately producing more clay to be translocated downward. The process accelerates by itself (Huang et al., 2012, Ch. 30, p. 30-16).
• Example (desertification): Destruction of vegetation cover → increased erosion → depletion of the upper horizon → death of remaining plants → further increased erosion (a degradation loop) (Huang et al., 2012, Ch. 38, p. 38-10).

Negative feedback (slows down, stabilizes):

• Example (organic matter accumulation): Increasing soil organic matter improves structure, water-holding capacity, and nutrient regime, stimulating plant growth and new organic inputs. However, microbial activity also increases, decomposing organic matter faster. Ultimately, the system may reach a steady state where input balances decomposition (Birkeland, 1984, p. 168; White, 2006, p. 100).

Feedbacks explain why soils often exhibit stable states (resilience), but upon crossing a threshold, they may sharply switch to another stable state (e.g., from forest to steppe, or from meadow to desert).

8.2.3. Heterogeneity and Self-organization

Modern research shows that even under seemingly identical factors, soils can exhibit significant spatial variability that cannot be explained by simple variation of one of the five factors (White, 2006, p. 89, Box 5.1). This is due to:

• Small but critical initial differences: Micro-relief irregularities, local centers of biological activity (burrows, root channels) can initiate nonlinear chains of processes that over time lead to noticeable differences ("butterfly effect" in soil formation) (Huang et al., 2012, Ch. 30, p. 30-22).
• Self-organization: Soil as an open thermodynamic system is capable of forming stable spatial patterns (e.g., tundra polygons, gilgai, desert stripes) that arise from internal interactions rather than being a direct reflection of external factors (Phillips, 1998, cited in White, 2006, p. 89).

8.3. Humans as a Full-Fledged Factor: Anthropogenesis

We mentioned humans in the organisms chapter, but in modern additions, humans deserve separate consideration. Today, human activity is no longer just an "external disturbance" but a full-fledged factor of soil formation (Huang et al., 2012, Ch. 38; Weil and Brady, 2017, Ch. 2, p. 79).

Scale of impact: More than half of the Earth's land surface is somehow affected by human activity (plowing, pastures, urbanization, deforestation, amelioration) (Huang et al., 2012, Ch. 38, p. 38-2; Weil and Brady, 2017, Ch. 1, p. 20).

New factors: Humans introduce new impacts into the system:

• Mechanical: plowing, terracing, construction, compaction.
• Chemical: fertilizer application, ameliorants (lime, gypsum), pollution (heavy metals, organic pollutants, acid rain, salts from irrigation).
• Biological: introduction of new plant and microorganism species, destruction of natural biota.
• Hydrological: drainage, irrigation, reservoir creation.

New "soils": As a result of these impacts, anthropogenic soils (Culturezems, Agrozems, Urbanozems, Technozems) are formed, whose properties differ radically from natural analogs (Huang et al., 2012, Ch. 38, p. 38-3; Scheffer et al., 2018, p. 352).

Metapedogenesis (after Yaalon and Yaron): A term introduced to describe soil formation processes under human influence, which can accelerate or slow natural processes, as well as create completely new ones (Huang et al., 2012, Ch. 38, p. 38-3). This extends the factorial model by adding a sixth factor—anthropogenic (h).

In the modern world, human impact often overshadows the influence of all other factors; therefore, when studying any soil, we must consider its land-use history.

8.4. Unifying Model: Soil as a Self-Organizing System

Ultimately, the modern understanding of soil formation can be summarized in several key points that complement the classical model:

1. Soil is an open system: It actively exchanges matter and energy with the surrounding environment (atmosphere, hydrosphere, biota). This creates conditions for self-organization and nonlinear dynamics (White, 2006, p. 89).

2. Multiple stable states: Soil can exist in different stable states under the same set of external factors (e.g., forest vs. meadow ecosystems on the same parent rock). Transitions between them occur abruptly, upon threshold crossing.

3. Hierarchical organization: Soil formation can be considered at different levels—from molecular to continental. Different mechanisms and patterns operate at each level (White, 2006, Box 5.1, p. 89).

4. Mutual influence of factors: Factors are not independent. For example, relief affects climate (altitudinal zonality), vegetation affects parent rock (biogenic weathering), and humans can alter climate (global warming). This creates a complex network of cause-and-effect relationships.

5. Soil memory: Soil "remembers" previous conditions (e.g., relict features of humid climate in modern arid regions). This phenomenon, called paleopedogenesis or relict features, often confuses the interpretation of modern factors (Scheffer et al., 2018, p. 378; Birkeland, 1984, p. 166).

Final Summary of the Entire Lecture:

We have traveled from Dokuchaev's simple yet brilliant idea of five factors, through Jenny's mathematical formalization, to a modern systemic view of soil formation. Along the way, we answered the main question: why can completely different soils be found nearby?

Because:

1. Climate (precipitation, temperature, their seasonality) creates different energy and moisture regimes.

2. Organisms (type of vegetation, composition of microfauna) transform mineral mass and shape the profile differently.

3. Relief (altitude, aspect, steepness, position) redistributes heat, moisture, and material, creating local centers of accumulation or erosion.

4. Parent rock (mineralogical, granulometric, chemical composition) sets the starting conditions—the "building material" for the soil.

5. Time (absolute and relative age) integrates the effects of all processes, allowing the soil to reach different degrees of development.

6. Modern factors (catenas, feedbacks, thresholds, humans) complicate the picture, making soil a self-organizing system capable of nonlinear and abrupt changes.

Soil is not just a product but an active participant in ecosystem processes, which both reflects the history of factors and itself shapes the environment for its further development. Understanding this complexity is the foundation for sound land use, soil conservation, and prediction of soil changes under global climatic and anthropogenic shifts.

In the following lectures, we will examine in detail specific soil processes (weathering, humification, eluviation, illuviation, etc.), which are the "tools" for implementing the action of these factors and lead to the formation of diverse soil profiles.