Thermal Regime of Soils

The Primary Source: The Sun

The absolute and practically only source of energy for the soil thermal regime is solar radiation (Mukha et al., 2003). The Earth’s internal heat, radioactive decay energy, or heat released during decomposition of organic matter are negligibly small compared to the energy flux we receive from the Sun, and at the scale of the surface soil horizon they can be ignored (Mukha et al., 2003).

However, the path of solar energy to the soil surface is not straightforward. Let us trace it step by step.

1. Solar constant and the Earth's atmosphere

Outside the Earth's atmosphere, the solar energy flux on a perpendicular surface is practically constant. This value is called the solar constant and amounts to about 1366 W/m² (Huang et al., 2012). This is the energy that could reach the Earth if it had no atmosphere.

But the atmosphere is a powerful filter. As solar rays pass through it, they interact with gas molecules, water vapour, aerosols, and clouds. As a result, part of the energy:

• Is absorbed (e.g., the ozone layer absorbs ultraviolet radiation; water vapour and carbon dioxide absorb infrared radiation).
• Is scattered (by air molecules and aerosols, which creates the blue sky and provides diffuse radiation).
• Is reflected back to space (especially by clouds).

In the end, only about 75% of the energy reaches the Earth’s surface on a clear day, and in cloudy regions even less, about 35–40% (Weil & Brady, 2017). The energy that reaches the surface is called total (global) radiation. It consists of two components:

• Direct radiation – rays coming directly from the solar disk.
• Diffuse radiation – rays that have changed direction in the atmosphere and arrive from the entire sky dome.

The proportion of direct and diffuse radiation depends on atmospheric conditions. On a clear day, direct radiation prevails; on a cloudy day, diffuse radiation prevails.

2. Interaction of radiation with the soil surface

Once it reaches the surface, total radiation is not entirely used for heating. Part of it is immediately reflected. This ability of the surface to reflect solar energy is, as we have already mentioned, called albedo.

Albedo is a key factor determining what proportion of energy goes toward heating and what proportion is lost immediately by reflection.

The factors that determine albedo are inseparably linked to the properties of the soil itself and vegetation, which illustrates our main thesis: heating depends not only on the sun.

• Soil colour. This is the most obvious factor. Soils rich in humus (e.g., chernozems) have a dark colour and an albedo of 8–14%. Light soils, such as grey soils or sands, reflect 25–40% or more (Mukha et al., 2003; Weil & Brady, 2017). The difference in absorbed energy between dark and light soils can reach 10–15% or more.

• Soil moisture. This is an interesting and important factor. Soil wetted with water visually becomes darker, and its albedo decreases. For example, the albedo of dry chernozem is about 14%, while that of wet chernozem drops to 8% (Mukha et al., 2003). Wet sand can have an albedo of 0.09–0.20 versus 0.35–0.45 for dry sand (Huang et al., 2012; Shukla, 2023). It would seem that wet soil should heat up more, but here the effect of latent heat (evaporation), which we discussed in the previous chapter, comes into play. Thus, moisture exerts a dual influence.

• Surface roughness. A smooth, even surface reflects more than an uneven, rough surface. For example, a freshly ploughed field has a lower albedo and heats up faster than a compacted one.

• Vegetation cover. The albedo of plants averages 15–25%, which is comparable to the albedo of many soils. However, vegetation fundamentally changes the radiation balance. It not only reflects part of the energy but also absorbs it for photosynthesis and also creates shade, protecting the soil surface from direct heating (Huang et al., 2012; White, 2006). Thus, even if the albedo of vegetation and soil are the same, the amount of energy reaching the soil surface itself will be radically different.

• Angle of incidence of sunlight (latitude and time of day). In the morning and evening hours, as well as in winter, the sun stands low above the horizon. Rays pass through a longer atmospheric path, are scattered more, and strike the surface at an acute angle. Consequently, the energy flux per unit area decreases, and heating proceeds more slowly (Weil & Brady, 2017).

Secondary Sources: Heat Exchange with the Atmosphere

In addition to absorbing direct and diffuse solar radiation, the soil exchanges energy with the air and other bodies. This flux, which we denoted as H in the balance equation, is a secondary source or sink of energy.

• Long‑wave radiation. Soil, like any heated body, emits energy in the infrared (thermal) range. This heat flux depends on the surface temperature and its properties. At night, when there is no sun, the soil can cool by giving off heat to the atmosphere. This phenomenon is called effective radiation.
• Atmospheric counter‑radiation. The atmosphere, in turn, also radiates heat back to the soil. This is especially important on cloudy nights, when clouds act as a "blanket," reflecting heat back and slowing cooling (Eash et al., 2016).

Summary on Energy Sources

Thus, we see that the amount of energy that actually enters the soil for heating is determined by a complex set of factors, not only by the intensity of sunlight. These factors can be divided into three main groups:

1. Atmospheric factors (cloudiness, air humidity, presence of aerosols), which regulate the transparency of the atmosphere to solar rays.

2. Geometric factors (latitude, time of year, time of day, slope angle), which determine the angle of incidence of rays.

3. Properties of the surface itself (colour, moisture, roughness, presence of vegetation), which determine what proportion of the energy that reaches the surface will be absorbed.

It is the third item that gives us, as soil scientists, the key to understanding why, under the same sunny weather, soils have different temperatures. It also opens opportunities for managing the thermal regime, which we will discuss later.

Connection with the Previous and Next Chapters

Now that we have dealt with the first half of the energy balance—where the energy comes from—we can move on to the second half: how the soil receives this energy.

In the next chapter, we will consider the fundamental thermal properties of soil. The first of these will be heat capacity—the soil's ability to accumulate heat. It is heat capacity that determines how strongly the soil temperature will rise upon absorbing a certain amount of energy. We will see that this property depends as strongly on soil composition (and, in particular, on water content) as albedo does.

Soil Components and Their Heat Capacity

Soil is a three‑phase system consisting of solid (mineral particles and organic matter), liquid (water), and gaseous (air) phases. Each of these phases has its own heat capacity, and the differences between them are enormous (Foth, 1990; Scheffer et al., 2018).

Let us look at the volumetric heat capacity values of the main soil components (approximate values at room temperature):

(Data from: Foth, 1990; Marshall et al., 1996; Weil & Brady, 2017; Scheffer et al., 2018)

Notice the enormous differences in these values! The heat capacity of water is about 4 times higher than that of mineral particles and almost 3500 times higher than that of air (Eash et al., 2016; Shukla, 2023). This means that to heat the same volume of water, one needs to spend 4 times more energy than to heat the same volume of mineral particles, and thousands of times more than to heat air.

Calculation of Soil Heat Capacity

Since soil is a mixture, its volumetric heat capacity is calculated as the weighted average sum of the heat capacities of all its components, taking into account their volume fractions (Shukla, 2023; Scheffer et al., 2018).

The calculation formula is as follows:

Where:

• C_v is the desired volumetric heat capacity of the soil.
• C_min, C_org, C_water, C_air are the volumetric heat capacities of the respective components.
• θ_min, θ_org, θ_water, θ_air are their volume fractions in the soil.

The contribution of air to soil heat capacity can often be neglected because of its negligibly small value.

Practical conclusion: Since the volume fraction of water in soil can vary greatly (from almost 0 to 50% or more), moisture is the main factor determining soil heat capacity (Huang et al., 2012; Shukla, 2023).

Example:

Consider two soils with the same volume and the same mineral composition, but different moisture content.

• Soil A (dry): contains 10% water (θ_water = 0.10).
• Soil B (wet): contains 30% water (θ_water = 0.30).

The heat capacity of dry soil A will be significantly lower than that of wet soil B. This means that under the same input of solar energy, dry soil A will heat up faster and to a higher temperature than wet soil B. This is why wet soils warm up more slowly in spring and are considered "cold," while dry sandy soils are considered "warm" (Weil & Brady, 2017).

Role of Organic Matter

Soil organic matter, despite its usually small volume (only a few percent), also contributes to heat capacity. Its heat capacity is higher than that of mineral particles, but lower than that of water. Therefore, soils rich in humus also have a somewhat higher heat capacity. However, the contribution of organic matter is generally not as great as that of water (Marshall et al., 1996).

Influence of Bulk Density

Volumetric heat capacity is also affected by bulk density (or dry bulk density) of the soil (Scheffer et al., 2018). The more densely packed the solid particles are in a unit volume, the greater the mass of mineral material, and therefore the higher the volumetric heat capacity. However, this effect is considerably weaker than the influence of moisture, and it is often treated as a secondary factor (Huang et al., 2012).

Summary on Heat Capacity

Thus, we have approached the answer to our main question from a new angle.

Why does soil temperature depend not only on sunlight? Because, even if they receive the same amount of solar energy, different soils will heat up differently. This depends on their heat capacity:

1. The main factor is moisture. Water has an anomalously high heat capacity. The more water in the soil, the more heat is required to warm it, and the slower it warms up.

2. Secondary factors are organic matter content and bulk density. They also affect heat capacity, but to a much lesser degree.

This property explains many everyday observations: why wet sand feels colder than dry sand, why soil cools after rain, and why wet clay soils remain cold for a long time in spring.

In the next chapter, we will consider another critical property that determines how quickly heat will spread through the soil profile. This property is thermal conductivity. We will see that high heat capacity slows heating, while high thermal conductivity, on the contrary, promotes rapid heat propagation downward.

Soil Components and Their Thermal Conductivity

As with heat capacity, the thermal conductivity of soil is an integral property that depends on the thermal conductivity of its individual components. But here the situation is radically different: the differences in thermal conductivity between components are even more striking than those in heat capacity (Foth, 1990; Marshall et al., 1996).

(Data from: Foth, 1990; Marshall et al., 1996; Weil & Brady, 2017; Scheffer et al., 2018)

Note:

1. Mineral particles, especially quartz, have a very high thermal conductivity. They are excellent heat conductors.

2. Water conducts heat less well than minerals, but much better than air.

3. Air is an excellent thermal insulator. Its thermal conductivity is more than 20 times lower than that of water and hundreds of times lower than that of minerals.

What Determines Soil Thermal Conductivity?

Since soil is a mixture of solid particles, water, and air, its overall thermal conductivity will depend on how these components are in contact with each other. Here two factors come to the fore:

1. Soil moisture.

This is undoubtedly the most important factor (Foth, 1990; Weil & Brady, 2017). In dry soil, the pores are filled with air. Heat is transferred mainly through contacts between solid particles, which, firstly, are few, and secondly, the air in the pores acts as an insulator, hindering heat transfer between particles. Therefore, the thermal conductivity of dry soil is very low.

When the soil becomes wet, water begins to fill the pores. It displaces air and, more importantly, creates water bridges between solid particles at their contact points. Water, having significantly higher thermal conductivity than air, becomes an additional channel for heat transfer. The thermal conductivity of soil increases sharply with increasing moisture (Marshall et al., 1996; Weil & Brady, 2017). This effect is especially strong during the transition from a very dry state to a moderately wet state.

2. Bulk density and structure.

The more densely packed the particles, the more contacts there are through which heat can be transferred. Therefore, compacted soil generally has higher thermal conductivity than loose soil (Scheffer et al., 2018). This is also because compacted soil has fewer pores filled with insulating air (Weil & Brady, 2017).

3. Mineralogical composition.

As can be seen from the table, quartz is an exceptionally good heat conductor. Therefore, sandy soils, consisting mainly of quartz, have, all else being equal, higher thermal conductivity than clay soils, where other minerals predominate.

4. Organic matter.

Organic matter has low thermal conductivity, so its accumulation in the upper horizons reduces the soil's ability to conduct heat (Marshall et al., 1996).

Relationship Between Heat Capacity and Thermal Conductivity

These two properties work in tandem and determine exactly how temperature will change over time and with depth. To describe the rate of temperature propagation, a combined characteristic is used — thermal diffusivity (or coefficient of thermal diffusivity).

Thermal diffusivity *(α) = λ / C_v*

Where:

• α is thermal diffusivity (m²/s or cm²/s).
• λ is thermal conductivity.
• C_v is volumetric heat capacity.

Thermal diffusivity indicates how quickly a temperature wave propagates in the soil. High thermal conductivity promotes rapid heat propagation, but high heat capacity, on the contrary, slows this process because much of the energy is spent on heating the soil itself (Marshall et al., 1996; Shukla, 2023).

Example: For water and quartz, the ratio of thermal conductivity to heat capacity is such that the thermal diffusivity of soil often reaches a maximum at some intermediate moisture content, not at full saturation. With further wetting, the increase in heat capacity begins to outpace the increase in thermal conductivity.

Summary on Thermal Conductivity

We have taken one more step closer to understanding the main question.

Why does soil temperature depend not only on sunlight? Because the ability of soil to conduct heat, like its ability to store heat, depends strongly on its composition and state, and therefore on the factors that we can observe and measure:

1. Moisture is the main controlling factor. Water not only increases heat capacity (slowing heating) but also greatly increases thermal conductivity, accelerating the transfer of heat from the surface downward.

2. Bulk density affects the number of contacts between particles and the volume of pores filled with insulating air.

3. Structure and mineralogical composition determine the baseline level of thermal conductivity.

In the next chapter, we will examine the consequence of these properties — temperature gradients. We will see exactly how the difference in thermal conductivity and heat capacity creates that "delayed" temperature profile with depth that we observe every day.

What is a Temperature Gradient?

Temperature gradient is a vector physical quantity that shows how quickly and in what direction temperature changes in space. In the simplest one‑dimensional case (with depth in the soil), the gradient is the change in temperature per unit length (e.g., per 1 cm or 1 m of depth). It is denoted as dT/dz, where T is temperature and z is depth (Eash et al., 2016; Huang et al., 2012).

According to Fourier's law (from the previous chapter), it is the magnitude of the temperature gradient that is the driving force for the heat flux. The steeper the gradient (the faster temperature changes with depth), the more intense the heat flux.

Types of Temperature Gradients in Soil

In the soil profile, we can distinguish two fundamentally different types of temperature distribution, which replace each other during the day and throughout the year.

1. Summer (daytime) gradient (or heating gradient).

During the warm season, especially on sunny days, the surface temperature is significantly higher than that of the underlying layers. Heat comes from above and spreads downward. In this case:

• Direction of heat flux: downward, from the surface to the subsoil.
• Temperature decreases with depth. The highest temperatures are at the surface, the lowest at some depth.
• The temperature gradient is directed downward. This means that with increasing depth, temperature decreases. This gradient is called positive or direct (Marshall et al., 1996; Weil & Brady, 2017).

2. Winter (night‑time) gradient (or cooling gradient).

In the cold season, as well as at night, the soil surface cools faster than the deeper layers. Heat begins to move from the bottom upward, to the surface, to compensate for losses due to radiation and air heating. In this case:

• Direction of heat flux: upward, from the subsoil to the surface.
• Temperature increases with depth. The lowest temperatures are at the surface, higher temperatures are at depth.
• The temperature gradient is directed upward. With increasing depth, temperature rises. This gradient is called negative or reverse (Marshall et al., 1996; Weil & Brady, 2017).

The transition from one type of gradient to the other occurs in the morning and evening hours, when surface and deep temperatures equalise. At that moment, the heat flux stops, and the gradient becomes close to zero (the so‑called isothermal state).

Temperature Waves: Diurnal and Annual

Since surface heating and cooling occur cyclically (day‑night alternation, seasons), periodic temperature fluctuations arise in the soil and propagate downward in the form of temperature waves (Marshall et al., 1996; Weil & Brady, 2017).

A diurnal temperature wave has a period of 24 hours. It is caused by daytime heating and night‑time cooling of the surface.

An annual temperature wave has a period of 365 days. It is caused by the change of seasons: summer heating and winter cooling.

Basic Laws of Temperature Wave Propagation

The propagation of a temperature wave in the soil obeys two fundamental laws, which are consequences of the combined action of heat capacity and thermal conductivity (Marshall et al., 1996; Shukla, 2023; Weil & Brady, 2017):

1. Attenuation of amplitude with depth.

The deeper we go into the soil, the smaller the temperature fluctuations around the mean value become. The maximum amplitude (the difference between the highest and lowest temperatures) is observed at the very surface. At a depth of 10–20 cm, the amplitude of diurnal fluctuations decreases several times, and at a depth of about 50 cm, diurnal fluctuations are practically completely damped (Eash et al., 2016; Weil & Brady, 2017). The annual wave, being much longer, penetrates much deeper—down to 10–15 metres (Marshall et al., 1996; Weil & Brady, 2017).

This happens because, along the path of heat propagation, the soil acts as a "filter": part of the energy is spent on heating the material itself (the effect of heat capacity), and part on overcoming resistance (limited thermal conductivity). As a result, the wave energy is gradually "dissipated."

2. Phase shift (lag) with depth.

The temperature wave not only attenuates, but also "lags" in phase. This means that the peak of maximum temperature (or minimum) at depth occurs later than at the surface. For example, if the maximum temperature at the surface is observed at 14:00, then at a depth of 20 cm it may occur only in the evening, and at a depth of 1 metre—with a delay of several days (Foth, 1990; Weil & Brady, 2017).

For the annual wave, this shift is even more noticeable: the highest temperatures at the surface occur in July, while at a depth of 2–3 metres they may be observed only in September or October (Huang et al., 2012; Marshall et al., 1996). This is because heat needs time to "pass" through the soil thickness. The speed of wave propagation depends on the thermal diffusivity of the soil: the higher it is, the faster the wave moves and the smaller the lag.

The Factor Determining Wave Propagation: Thermal Diffusivity

The speed of propagation and attenuation of the temperature wave are determined by the composite property we have already mentioned—thermal diffusivity (α). Recall:

Where:

• λ is thermal conductivity (ability to conduct heat).
• C_v is volumetric heat capacity (ability to store heat).

The higher the thermal diffusivity (for example, in dense, well heat‑conducting soil), the faster the wave propagates and the deeper it penetrates. The lower the thermal diffusivity (e.g., in loose, dry soil with low thermal conductivity), the slower the wave travels and the faster it attenuates (Marshall et al., 1996; Shukla, 2023).

Example: Sandy soil, having high thermal conductivity, conducts the thermal wave faster and deeper than, for example, wet clay soil. However, the high heat capacity of wet clay soil slows its heating, but also slows its cooling. This leads to the fact that in clay soil, diurnal temperature fluctuations attenuate faster, but annual fluctuations may be more smoothed.

Summary on Temperature Gradients

We have completed our review of the key physical properties that determine the soil thermal regime. The answer to our main question now sounds complete and well‑argued.

Why does soil temperature depend not only on sunlight?

Because:

1. Energy input depends on the surface albedo, which is determined by colour, moisture, structure, and vegetation cover (Chapter 2).

2. Energy storage depends on soil heat capacity, which is primarily determined by water content (Chapter 3).

3. Energy propagation depends on thermal conductivity, which is also strongly influenced by soil moisture and density (Chapter 4).

4. Temporal and spatial dynamics (gradients, waves, lag) are determined by thermal diffusivity—the ratio of thermal conductivity to heat capacity (Chapter 5).

Thus, the same solar energy is transformed differently into the thermal regime of different soils. This knowledge gives us the key not only to understanding nature but also to managing the thermal state of the soil.

In the next chapter, we will pay special attention to the role of water, since, as we have repeatedly seen, it is the main "conductor" of the thermal regime.

1. Water and Albedo (Energy Input)

As we have already said, water changes the reflective ability of the surface. Wet soil is always darker than dry soil, so its albedo decreases (Huang et al., 2012; Mukha et al., 2003). This means that wet soil absorbs more solar energy than dry soil. This would seem to favour heating. However, as we will see below, other effects of water more than offset this.

2. Water and Heat Capacity (Energy Storage)

This is perhaps the most well‑known and significant effect of water. As we recall from Chapter 3, the volumetric heat capacity of water is about 4 times higher than that of mineral particles and thousands of times higher than that of air (Eash et al., 2016; Scheffer et al., 2018).

This means that:

• Wet soil requires significantly more energy to warm up than dry soil. Therefore, wet soils warm up slowly in spring and are considered "cold."
• Wet soil cools down more slowly. The heat stored during the day will be released to the atmosphere at night longer than from dry soil. This creates a smoothing effect, reducing diurnal and seasonal temperature fluctuations (Weil & Brady, 2017; White, 2006).

Practical conclusion: If we want the soil to warm up faster in spring, we should promote the removal of excess water (drainage, ridging). If we want to protect the soil from overheating in summer, we retain moisture (mulching).

3. Water and Thermal Conductivity (Heat Transfer)

Water also drastically changes the thermal conductivity of soil. In dry soil, the air in the pores acts as a thermal insulator, and heat is transferred mainly through rare contacts between particles. When the soil becomes wet, water:

• Fills the pores, displacing air with its extremely low thermal conductivity.
• Creates water bridges between solid particles, creating new, continuous pathways for heat transfer (Foth, 1990; Marshall et al., 1996).

As a result, soil thermal conductivity sharply increases with increasing moisture, especially at the initial stage of wetting (Weil & Brady, 2017). This means that heat absorbed at the surface is transferred faster and more efficiently to deeper layers. Wet soil not only heats up more slowly but also conducts this heat downward faster.

4. Water and Thermal Diffusivity (Speed of Heat Propagation)

The combined effect of water on heat capacity and thermal conductivity leads to a complex, nonlinear effect on thermal diffusivity (α = λ / C_v).

At low moisture, adding a small amount of water significantly increases thermal conductivity while weakly affecting heat capacity. As a result, thermal diffusivity increases—the thermal wave propagates faster. However, with further increases in moisture, the growth of heat capacity begins to outpace the growth of thermal conductivity, and thermal diffusivity may decrease (Marshall et al., 1996; Shukla, 2023).

Thus, the dependence of thermal diffusivity on moisture has a bell‑shaped form, with a maximum at some intermediate moisture content, which is often close to field capacity (Weil & Brady, 2017).

5. Water and Phase Transitions (Latent Heat)

This is perhaps the most powerful and dramatic effect of water on the thermal regime. Water in soil can exist in three phases: liquid, solid (ice), and gaseous (water vapour). The transition from one phase to another is accompanied by the absorption or release of a huge amount of energy—latent heat (Eash et al., 2016; Weil & Brady, 2017).

• Evaporation (liquid → gas). Evaporating 1 gram of water requires about 2.5 kJ of energy. This energy is taken from the soil, leading to its cooling. This is why wet soil, from which water actively evaporates, is always colder than dry soil, even with the same solar energy input. This effect is especially important in hot weather (Eash et al., 2016; Weil & Brady, 2017).

• Condensation (gas → liquid). When water vapour condenses in the soil (for example, at night when the surface cools), the same amount of heat is released. This contributes to warming the soil. Condensation is an important but often underestimated source of heat in the upper soil layers (Foth, 1990).

• Freezing (liquid → solid). When water freezes, about 0.33 kJ per 1 gram is released (heat of crystallisation). This slows down soil freezing because the released heat counteracts further cooling (Weil & Brady, 2017). This is why wet soils freeze more slowly than dry ones.

• Melting (solid → liquid). This process requires the absorption of the same amount of heat. Therefore, spring thawing of soil is an energy‑intensive process that slows its warming (Foth, 1990).

6. Water and Heat Transport (Convection)

In addition to thermal conductivity, heat can be transported in soil with moving water—this is convective transport (Huang et al., 2012; White, 2006). Infiltrating rainwater, which may be warmer or colder than the soil, carries thermal energy with it. The same happens when soil moisture moves in capillaries. Although this mechanism is less significant for the steady‑state regime, it can be very important during periods of active precipitation or irrigation.

Summary on the Role of Water

Thus, water affects the thermal regime of soil in all directions:

1. Increases energy absorption (lowers albedo).

2. Increases heat capacity (slows heating and cooling).

3. Increases thermal conductivity (accelerates heat transfer downward).

4. Nonlinearly affects thermal diffusivity, passing through a maximum.

5. Changes the heat balance through phase transitions, causing cooling upon evaporation, warming upon condensation, slowing freezing upon freezing, and slowing thawing upon melting.

6. Creates convective heat transport with moving water.

Water is the factor that makes the soil thermal regime complex, dynamic, and, most importantly, manageable. By understanding these mechanisms, we can answer why, under the same weather conditions, soils have different temperatures, and we can purposefully influence this process.

Connection with the Previous and Next Chapters

We have examined the physical foundations of the thermal regime, from energy input to its propagation, and we have seen that water is the central link. Now we move on to the next important aspect—the influence of soil structure. How exactly do aggregate state, bulk density, and the presence of macro‑ and micropores affect thermal properties and the distribution of water within them? This is the topic of the next chapter.

How Does Structure Affect the Thermal Regime?

The influence of structure on the thermal regime manifests through several key mechanisms, which we will consider in order.

1. Bulk density and contacts between particles

Bulk density (or dry bulk density) is essentially an indicator of how tightly solid particles are packed in a unit volume of soil. This parameter directly affects thermal conductivity (Marshall et al., 1996; Weil & Brady, 2017).

• The denser the soil, the more contact points there are between mineral particles. Heat in the solid phase is transmitted precisely through these contacts (through so‑called "contact thermal conductivity"). The more such contacts, the more efficient the heat transfer.
• The denser the soil, the fewer large pores filled with air. Air, as we recall, is an excellent thermal insulator. Reducing its volume reduces the insulating effect and increases thermal conductivity (Weil & Brady, 2017).

Thus, compacted, well‑aggregated soil generally has higher thermal conductivity than loose, structureless soil. However, it is also important to remember heat capacity. Compaction increases the mass of solid material per unit volume, which slightly increases heat capacity as well, but this effect is usually much weaker than the effect on thermal conductivity.

2. Pore space: macro‑ and micropores

Soil structure creates a hierarchy of pores of different sizes. This division into macropores (large, > 50–75 µm) and micropores (small, < 30 µm) has enormous significance for the thermal regime (Foth, 1990; Scheffer et al., 2018).

• Macropores are large voids between aggregates, cracks, wormholes, and root channels. They are generally filled with air, unless the soil is saturated. Air in macropores acts as a thermal insulator, slowing heat transfer between aggregates (Weil & Brady, 2017).
• Micropores are small pores inside aggregates and between fine particles. It is in these that the bulk of capillary water is retained. This water, having high thermal conductivity, creates continuous water bridges linking particles within the aggregate and promoting efficient heat transfer (Marshall et al., 1996).

The ratio of macro‑ to micropores determines how strongly the effect of water manifests. In a well‑structured soil with many macropores, water will accumulate mainly in micropores, not completely blocking the insulating air in large pores. This can lead to faster warming of the upper layers, because heat will be transferred within aggregates through water, but insulation between aggregates will be maintained.

3. Role of aggregation

Aggregates are structural units consisting of many primary particles bound together. Aggregation creates dual porosity: within aggregates (micropores) and between them (macropores). This leads to heterogeneity of thermal properties even within a single horizon (Foth, 1990; Weil & Brady, 2017).

• Inside the aggregate, micropores are often filled with water, which ensures high thermal conductivity inside the aggregate itself.
• Between aggregates, macropores are filled with air, creating insulating interlayers.

As a result, the thermal wave propagates differently through aggregates and through inter‑aggregate spaces. This leads to a more complex and smoothed temperature pattern compared to homogeneous, structureless soil. In addition, aggregates act as a kind of "buffers"—they heat and cool slowly, smoothing out sharp diurnal temperature fluctuations at the surface (Weil & Brady, 2017).

4. Influence of structure on water regime and phase transitions

Structure not only determines where water is located, but also how it moves. A well‑developed aggregated structure with macropores ensures rapid drainage of excess moisture in spring. This, in turn, accelerates soil warming because it reduces heat capacity and evaporation (Foth, 1990; Weil & Brady, 2017). In contrast, structureless, compacted soil with a predominance of micropores will retain water longer, warm up more slowly, and cool more strongly due to evaporation.

In addition, structure affects the depth and nature of freezing. Large macropores filled with air are poor heat conductors, so soil with good macrostructure freezes more slowly and to a shallower depth than dense, structureless soil (Weil & Brady, 2017). The presence of aggregates also creates additional interfaces where water vapour can condense, releasing heat and slowing freezing.

Summary on the Influence of Structure

Soil structure is not just a "skeleton" on which the soil profile is built. It is an active element that determines exactly how water and heat interact in the soil. The influence of structure on the thermal regime is multifaceted:

1. Determines bulk density and, consequently, the number of contacts between particles, affecting thermal conductivity.

2. Creates a hierarchy of pores (macro‑ and micropores), which determines where water is and where air is, and accordingly forms heterogeneity of thermal properties.

3. Mediates the effect of water on the thermal regime, determining its distribution and drainage rate.

4. Smooths temperature fluctuations, especially in the upper horizons, thanks to the presence of aggregates acting as thermal buffers.

Thus, structure is the "interface" that links together all the factors we have considered (albedo, heat capacity, thermal conductivity, water). By changing the structure (e.g., through tillage, application of organic fertilisers, creation of plant cover), we can influence the thermal regime indirectly, through redistribution of water and changes in physical properties.

In the next, concluding chapter, we will consider how surface factors—mulch, vegetation cover, and snow cover—affect the thermal regime, bringing our understanding to a practical level.

1. Organic Mulch (Plant Residues)

Organic mulch is a layer of plant residues (straw, leaves, grass, bark, sawdust) lying on the soil surface. This is perhaps the most powerful and multifaceted factor in managing the thermal regime (Weil & Brady, 2017; White, 2006).

The influence of mulch on the thermal regime manifests through several mechanisms, which can act in opposite directions, but together give a characteristic overall effect.

Effects of mulch:

1. Change in albedo. Dry plant mulch generally has a higher albedo (reflectivity) than dark, wet soil. This is especially noticeable for light straw. This results in more solar energy being reflected back to the atmosphere, and the soil surface under the mulch receives less energy for heating (Eash et al., 2016; Weil & Brady, 2017).

2. Thermal insulation. The mulch itself, especially if it consists of coarse particles (straw, bark), contains a lot of air between the fibres. Air, as we know, is an excellent thermal insulator. Therefore, the mulch layer slows heat transfer both from the soil to the atmosphere (at night) and from the atmosphere to the soil (during the day). This leads to smoothing of diurnal temperature fluctuations (Marshall et al., 1996; Weil & Brady, 2017).

3. Reduction of evaporation. Mulch physically blocks the path of water vapour leaving the soil. This means that soil moisture under mulch is maintained higher than on an open surface (Eash et al., 2016; Weil & Brady, 2017). And higher moisture, as we recall, increases heat capacity and therefore slows heating.

Net effect: In most cases, especially in cool climates, organic mulch lowers the daytime surface temperature of the soil, but may raise it at night, acting as a "blanket" that retains heat (Weil & Brady, 2017). As a result, diurnal temperature fluctuations become less sharp. This makes mulch an ideal means of protection against overheating in hot climates, but may slow soil warming in spring in cold regions (White, 2006).

2. Vegetation Cover

Living vegetation is an even more complex and effective regulator of the thermal regime than dead mulch. Plants act in several directions simultaneously (Huang et al., 2012; Weil & Brady, 2017).

1. Shading. The most obvious action. Green foliage creates shade that intercepts most of the direct solar radiation, preventing it from reaching the soil surface. This is especially effective under closed forest canopies or dense crop stands. As a result, soil under vegetation is significantly cooler than open soil (Eash et al., 2016).

2. Change in albedo. The albedo of vegetation (15–25%) is often higher than that of dark, wet soil, which also contributes to reflecting part of the energy (Huang et al., 2012).

3. Transpiration. This is perhaps the most powerful cooling mechanism. Plants actively evaporate water through leaves (transpiration). This process consumes a huge amount of energy (latent heat), which is taken from the environment, including the soil and the near‑surface air layer. Transpiration effectively cools both the plant itself and the soil beneath it (Weil & Brady, 2017; White, 2006).

4. Aerodynamic resistance. Vegetation cover creates a rough surface that reduces wind speed near the soil surface. This reduces turbulent heat exchange and evaporation from the soil surface itself, which also helps retain heat at night (Marshall et al., 1996).

Net effect: Vegetation cover is a powerful cooling factor during the daytime and a smoothing factor for diurnal and seasonal fluctuations. Soil under forest or dense grass is always cooler and has a more stable temperature than open soil (Weil & Brady, 2017).

3. Snow Cover

Snow is a unique surface factor that drastically changes the thermal regime during the cold season. Its influence is determined by its outstanding thermal insulation properties (Eash et al., 2016; Weil & Brady, 2017).

1. Extremely high albedo. Freshly fallen snow reflects up to 80–90% of solar radiation. This means that the soil under snow practically does not receive solar energy and does not warm up even on sunny winter days (Huang et al., 2012; Mukha et al., 2003).

2. Exceptionally low thermal conductivity. Snow cover, especially if loose, consists mainly of air trapped between snowflakes. The thermal conductivity of snow is very low. This means that snow cover acts as a powerful thermal insulator that prevents heat loss from the soil to the cold atmosphere (Marshall et al., 1996; Weil & Brady, 2017).

3. Smoothing temperature fluctuations. Thanks to its insulating properties, snow cover protects the soil from severe freezing. Under a thick layer of snow, the soil temperature can remain close to 0 °C even during severe frosts. This is especially important for winter crops, which can survive the winter only under reliable snow cover.

Net effect: Snow cover is the "winter blanket" for the soil. It prevents deep freezing, smooths winter temperature fluctuations, and, most importantly, contributes to moisture accumulation in spring when it melts (Weil & Brady, 2017; White, 2006).

Links Between Surface Factors

It is important to understand that these three types of surface conditions (mulch, vegetation, snow) often act together and may replace each other during the year. For example, in autumn, leaf litter creates a mulch layer; in winter, it is covered by snow; and in spring, active soil life begins under this layer. In agricultural landscapes, mulch from plant residues on fields (e.g., under no‑till) simultaneously performs the functions of erosion protection, moisture conservation, and thermal regime regulation (Weil & Brady, 2017).

Summary on the Role of the Surface

We have completed our journey from the general to the specific. We have considered all the key factors that determine the soil thermal regime. Now we can give a final and complete answer to the main question of the lecture:

Why does soil temperature depend not only on sunlight?

Because, on the path of solar energy to the soil and within it, a complex system of intermediaries operates. Soil temperature is the result of the integrated influence of the following factors:

1. Intensity of solar radiation (depends on latitude, time of year and day, cloudiness).

2. Surface properties (albedo, presence of mulch, vegetation or snow cover).

3. Soil water regime (moisture determines heat capacity, thermal conductivity, and energy expenditure on evaporation).

4. Soil structure and density (determine pore space, contacts between particles, and, accordingly, thermal conductivity).

5. Properties of the soil itself (mineralogical composition, organic matter content).

All these factors are interrelated. By changing one of them (for example, creating a mulching layer or regulating the water regime), we influence all the others and, ultimately, the thermal regime of the soil as a whole.

Understanding this system is the key to competent management of the thermal regime in agriculture, ecology, and landscape design.