What is Soil?

The soil is at the interface between the atmosphere and lithosphere (the mantle of rocks making up the Earth’s crust). It also has an interface with the hydrosphere, i.e. the sphere describing surface water, ground water, and oceans. The soil sustains the growth of many plants and animals, and so forms part of the biosphere. A combination of physical, chemical and biotic forces acts on organic and weathered rock fragments to produce soils with a porous fabric that contain water and air (pedosphere). Soil is as a natural body of mineral and organic material that is formed in response to many environmental factors and processes acting on and changing soil permanently.

Figure 1. Compartments of a landscape.

Because soil is important for cultivation and agricultural production, soil fertility and productivity are important issues to address. Detailed pedological knowledge is useful for land evaluation purposes, i.e., the classification in fertile productive soils and less valuable soils. Soils are an integral part of landscapes and the knowledge of the distribution of different soils helps to preserve a high standard in environmental quality.


The interpretation, mapping, and classification of soils based on soil genesis.

A pedon is the smallest volume that can be recognized as a soil individual. It has three dimensions and its area ranges from 1 to 10 square meters, depending on the variability in the horizons. The shape of the pedon is roughly hexagonal. A soil volume that consists of more than one pedon is termed a polypedon.

Peds are soil aggregates which are recognizable in the field because they are separated by voids and natural planes of weakness. They should persist through cycles of wetting and drying, as distinct from the less permanent aggregates.

The process of soil formation as the integral result of the combination of soil forming factors (climate, parent material, relief, organisms, time).

Soil classification:
The grouping of different soils into classes on the basis of one or more criteria. A class is a group of individuals or units similar in selected properties and distinguished from all other classes by differences in these properties. U.S. Soil Taxonomy : Classification of soils developed by the USDA Soil Survey Staff with the support of and contributions from land-grant university participants in the National Cooperative Soil Survey and colleagues in other countries. It was formally released and published by the USDA in 1975 (Soil Survey Staff, 1975) and is updated on a continuous basis (Soil Survey Staff, 1994).

Soil Survey Staff. 1975. Soil Taxonomy. U.S. Dept. Agric. Handbook 436.

U.S. Govt. Printing Office, Washington D.C. Soil Survey Staff. 1994. Keys to Soil Taxonomy USDA – Soil Conservation Service. 6th ed., Washington D.C.


SSSA: Glossary of Soil Science Terms
Landscaping and Soil Biology Glossery

Soil Profiles

The vertical dimension, exposed by excavation from the surface to the parent material, constitutes the soil profile. Layers in the soil, distinguished on the basis of color, soil texture, soil structure, and other visible properties, are called horizons. Several horizons make up the soil profile. Next, a description of a soil profile developed in a temperate humid environment is given showing all of the principle soil horizons (Figure 2.): The upper layer, from which materials are generally washed downwards, is described as eluvial. Lower layers in which these materials accumulate are called illuvial. The eluvial horizon is denoted by the notation E and an illuvial horizon by the notation B. The B and E horizon plus the top horizon denoted by an A comprise the solum. The unconsolidated parent material in which the soil is formed is denoted by a C. If unweathered rock (consolidated material) exists below the parent material it is labelled with a R. Organic litter on the surface, which is not incorporated in the soil is designed the O layer.

Fig. 2. A hypothetical soil profile showing all of the principle soil horizons.

Components of the Soil Soils are made of four main components, i.e.

mineral matter (40 – 60 %),
soil water (20 – 50 %),
soil air (0 – 40 %) and
organic material (small percentage).

The relative proportions of the four major components may vary widely, but generally lie within the ranges listed.

Concepts of Soil Genesis / Classification

Soil Classification

Classification of natural phenomena like soils is typically done for the following purposes:

  • Organize knowledge about the subject to enable investigation and communication to be both logical and comprehensive (structure/organization for scientific/technical development).
  • Provide a framework for establishing relationships among soils and their environment that leads to advancement of theoretical and experimental aspects of soil and related science (academic focus).
  • Establish groupings of soils for which useful and reliable interpretations can be made (utilitarian focus). For example: optimal use(s), hazard/limitation/remediation assessment, potential productivity, framework for technology transfer/information dissemination.

There are two different classification approaches, i.e., technical vs. natural classification. Technical classifications are designed for specific applied purposes (e.g., soil engineering classifications, based largely on physical properties), whereas natural classifications attempt to organize the divisions of soils from a more holistic appraisal of soil attributes.

The general principles of natural soil classification systems are:

  • A natural system of classification should express general or universal relationships that exist in nature. One should be able to understand, remember, generalize, or predict from information obtained.
  • The scheme should be based on characteristics or attributes of things classified as related to their genesis. It should place similar things together on the basis of their properties.
  • It is technically impossible to use all of the properties of soils to classify them. Judgment based on existing knowledge must be used to determine which properties are most important.

Historical Perspective

One of the earliest land evaluation systems that incorporated a soil classification was established during the Vao dynasty (2357-2261 B.C.) in China. Soils were graded into nine classes, based on their productivity. It has been suggested that property taxes were based on the size of the individual land holding and soil productivity. In former times (< 1600 A.C.), soil was solely considered as a medium for plant growth. Knowledge of soil behavior and crop growth was passed from generation to generation gained by observation. For example, in the Middle Ages it was well know that manure applied to soils improved crop growth. For instance, the 'Plaggen cultivation' was practiced for a long time in Europe, which left 'Plaggen soils': The top of grassland was peeled off and used as litter in the stables. This material mixed with manure was applied to arable land to improve crop production. In 1840, the German chemist Justus von Liebig initiated a revolution in soil science and agriculture. He proved that plants assimilate mineral nutrients from the soil and proposed the use of mineral fertilizers to fortify deficient soils. Crop production was increased tremendously using mineral fertilizers. Another effect was the shift from extensive to intensive techniques in agriculture, which influenced soils. Thaer (1853) published a classification that combined texture as a primary subdivision with further subdivisions based on agricultural suitability and productivity.

Several classifications based largely on geologic origin of soil material were also proposed in the 19th century (Fallou, 1862; Richtofen, 1886). From the 1660s onwards, various members of The Royal Society of London proposed schemes of soil classification that incorporated elements of a natural or scientific approach in their criteria. From this period on, the disciplines of agricultural chemistry (with a strong focus on soil fertility), geography, and geology provided a broad but somewhat fragmented background from which pedology emerged as a separate discipline in the late 19th century more or less independently in Russia (Dokuchaev and colleagues) and in the United States (Hilgard and colleagues).

In 1883, Dokuchaev carried out a comprehensive field study in Russia, where he described the occurrence of different soils thoroughly using soil morphologic features. Due to his observations in the field he hypothesized that different environmental conditions result in the development of different soils. He defined soil as an independent natural evolutionary body formed under the influence of five factors, of which he considered vegetation and climate the most important. Dokuchaev is generally credited with formalizing the concept of the ‘five soil forming factors’, which provides a scheme for study of soils as natural phenomena. The soil classification developed by Dokuchaev and his colleagues (Glinka, Neustruyev) was based on the soil forming factors -> soil forming processes -> and diagnostic horizons / soil properties. The focus in his soil classification approach was on soil genesis, therefore the classification system is called ‘genetic’.

In the United States, Hilgard (1892) emphasized the relationship between soils and climate, which is known as the climatic zone concept. Coffey (1912), produced the first soil classification system for the United States based on the soil genesis principles of Dokuchaev and Glinka. Marbut (1951) introduced the concepts of Coffey into soil survey programs in the U.S. carried out by the USDA (United States Department of Agriculture). Between 1912 and the 1960 the soil classification in the U.S. used a genetic approach. Jenny (1941) put together a detailed description of the five soil forming factors responsible for the development of different soils. In 1959, Simonson stressed that many genetic processes are simultaneously and / or sequentially active in any soil. Hence, a soil classification based on principles of soil genesis would not be favorable. Other soil scientists in the U.S., like Smith (director of Soil Survey Investigations for the USDA Soil Conservation Service) agreed, that soil genesis is very important for soil classification, but that genesis itself cannot be used as a basis for soil classification because the genetic processes can rarely be quantified or actually observed in the field.

Coffey G.N., 1912. A Study of the Soils of the United States. U.S. Dept. of Agriculture Bur., Soils Bull. 85. U.S. Govt. Printing Office, Washington D.C.

Dokuchaev V.V., 1883. Russian Chernozems (Russkii Chernozems). Israel Prog. Sci. Trans., Jerusalem, 1967. Transl. from Russian by N. Kaner. Available form U.S. Dept. of Commerce, Springfield, VA.

Fallou F.A., 1862. Pedologie oder allgemeine und besondere Bodenkunde, Dresden, Germany.

Hilgard E.W., 1892. A Report on the Relations of Soil to Climate. U.S. Dept. of Agriculture. Weather Bull. 3: 1-59.

Jenny H., 1941. Factors of Soil Formation. McGraw-Hill, New York.

Liebig J. von, 1840. Chemistry in its Application to Agriculture and Physiologie. Playfaeir, Taylor and Waton, London.

Marbut C.F., 1951. Soils: Their Genesis and Classification. A Memorial Volume of Lectures given in the Graduate School of the United States Dept. of Agriculture in 1928. Publ. 1951 by Soil Sci. Soc. Am., Madison, WI.

Richtofen F.F. von, 1886. Fuehrer fuer Forschungsreisende. Berlin.

Simonson R.W., 1959. Outline of a Generalized Theory of Soil Genesis. Soil Sci. Soc. Am. Proc. 23: 152-156.

Present Perspective in the U.S.

The soil classification based on external environmental factors and an assumed genesis as differentiating characteristics led to dissatisfaction among U.S. soil scientists in the 1950’s and 1960’s. It was stressed that there is a lot of uncertainty involved dealing with soil forming factors and changing environmental conditions in a landscape. Many processes go on in any soil, often offsetting one another. Hence, it is often difficult to identify the processes in soils because most soils are polygenetic. Furthermore, it is difficult to assess the relative importance of each soil forming factor contributing to the development of a distinct soil class. There was much concern of not being able to classify certain soils with adequate agreement because of uncertainties or disagreement concerning their genesis. The knowledge about processes was and still is limited, because we do not completely understand soil processes. It became obvious that a more ‘objective’ soil classification scheme should be developed to replace the ‘subjective’ one.

As a consequence, a completely new system of soil classification was developed: The U.S. Soil Taxonomy. It is termed a ‘natural’ soil classification system, which attempts to organize the division of soils from a more holistic appraisal of soil attributes. The pedon is used as the sampling unit for soil classification and mapping in schemes employing the U.S. Soil Taxonomy. Diagnostic horizons are used to define different soil taxa. Since that time, the Keys to Soil Taxonomy have undergone six published revisions.

The principles of Soil Taxonomy are:

  • Classify soils on basis of properties
  • Soil properties should be readily observable and / or measurable
  • Soil properties should either affect soil genesis or result from soil genesis

Soil processes are not refused in the Soil Taxonomy but the focus is on soil properties rather than on processes. The rationale behind this is that processes are important to produce soil properties and that those are used for classification, i.e., genesis ‘lies behind’ the soil taxa. Soil processes provide a framework for understanding the Soil Taxonomy. For example, the soil forming process of decomposition and humification forms an A horizons. The soil properties recognizable in the field are a dark color of the A horizon with a granular soil structure. The soil properties selected should be observable or measurable, though instruments may be required for observation or measurement. Properties that can be measured quantitatively are to be preferred to those that can be determined only qualitatively.

Other Perspectives

There are many diverse approaches to soil classification in different countries. Most schemes have in common a hierarchical structure, i.e. a structure in which individuals are collected in small groups, the small groups belonging to larger groups, a.s.o. The criteria to be used in establishing the different categories are:

  • The inferred genesis of the soil, on the assumption that the soil is in equilibrium with its environment and reflects the influence of prevailing soil forming factors, or:
  • The soil properties being the outward expression of the processes which shaped the soil.

No matter how divergent current soil classifications are in principle, language and format, two main levels of generalization stand out:

  • A lower level encompassing the surveyor’s series, which are narrow classes in harmony with the landscape.
  • A higher level of broad classes (conceptual classes) with somewhat overlapping boundaries, each of which embodies a central, usually pedogenetic concept.

Examples for modern soil classifications used in different countries:

FAO Russia United Kingdom Germany

Soil Formation

In the book by Jenny (1941) ‘Factors of Soil Formation’ it was presented an hypothesis that drew together many of the current ideas on soil formation, the inspiration for which was owned much to the earlier studies of Dokuchaev and the Russian school. The hypothesis was that soil is formed as a result of the interaction of many factors, the most important of which are:

  • Climate (cl)
  • Organisms (o)
  • Relief (r)
  • Parent Material (p)
  • Time (t)

Jenny’s approach was to consider these soil forming factors as control variables, independent of the soil as it evolved, and also independent, but not necessarily so of each other. He then attempted to define the relationship between any soil property ‘s‘ and the most important soil forming factors by a function of the form:

s = f (cl, o, r, p, t, ……)


The dots indicate that factors of lesser importance such as mineral accession from the atmosphere, or fire, might need to be taken into account. Equation 1 assumes that there is a causal relationship between s and the soil forming factors. Jenny (1980) redefined the soil forming factors as ‘state’ variables and included ecosystem properties, vegetation and animal properties, as well as soil properties. Parent material and relief define the initial state for soil development, climate and organisms determine the rate at which chemical and biological reactions occur in the soil (the pedogenic processes), and time measures the extent to which a reaction will have proceeded. There is a logical progression: of environment (i.e. the soil forming factors) -> processes -> soil properties underlying the soil formation.

Figure 1. Relationship between soil forming factors, processes, and soil properties.

To simplify the application of equation 1, it has been practice to solve it for changes in a soil property s when only one of the control variable (e.g. climate) varies, the others being constant or nearly so. The relationship is then called a climo-function (climate = control variable):

s = f (cl) o, r, p, t, …


and the range of soils formed is called a climosequence. Biosequences, toposequences, lithosequences, and chronosequences of soils have been recognized in various parts of the world. The term topo-sequence is synonymous with Milne’s catena concept (Milne, 1935). Indeed, the main virtue of Jenny’s attempt to quantify the relationship between soil properties and soil forming factors lies not in the prediction of exact values of s at a particular site, but rather in identifying trends in properties and soil groups that are associated with readily observable changes in climate, parent material, etc.

Jenny H., 1941. Factors of Soil Formation. McGraw-Hill, New York.

Jenny H., 1980. The Soil Resource, Origin and Behaviour, Springer-Verlag, New York.


Climate involves both local (microclimatic) and global (macroclimatic) considerations. The key components of climate in soil formation are moisture and temperature.Soil moisture depends on several factors:

  • The form and intensity of precipitation (water, snow, sleet)
  • Its seasonal variability
  • The transpiration and evaporation rate
  • Slope
  • Aspect
  • Depth of soil profile
  • Soil texture / permeability of the parent material

A method of defining the soil moisture regime of the soil is using water balance calculations. Such a calculation is based on the measurement of rainfall distribution, a calculation of the potential evapotranspiration, and an assessment of surface runoff and infiltration :

Water balance equation:

Inflow = Outflow +/- storage within the system

P = ET + SR + I +/- S



P: Precipitation (mm)
ET: Evapotranspiration (mm)
SR: Surface runoff (mm) (may include interflow)
I: Infiltration (mm)
S: Soil moisture storage (mm)

In the U.S. precipitation amounts (daily, hourly) and intensities (15-minute data) are collected at about 7000 weather stations. The potential evapotranspiration can be calculated by empirical equations (e.g. Thornthwaite) or by physically-based equations (e.g. Penman Monteith) (Maidment, 1992). The empirical Thornthwaite equation calculates the evapotranspiration in dependence of air temperature, whereas the Penman-Monteith equation is the currently most advanced resistance-based model of evapotranspiration. This equation allows the calculation of evapotranspiration form meteorological variables and resistances, which are related to the stomatal and aerodynamic characteristics of the crop. Infiltration and surface runoff can be calculated by empirical equations such as the Curve Number Method (Soil Conservation Service – USDA, 1985). This is a simple method, which calculates infiltration and surface runoff using land use and hydrologic soil groups to derive Curve Numbers, precipitation amount, initial abstraction values, and potential maximum retention of a soil. There are many other complex simulation models such as SWAT (Soil and Water Assessment Tool) (Arnold et al., 1993), WEPP (Water Erosion Prediction Project) (USDA-ARS, 1995), or OPUS (Smith, 1992), which calculate infiltration, surface runoff, and soil moisture.

The primary topographic attributes slope and aspect have a major impact on soil moisture. This was first expressed by Beven et al. (1979) in form of the compound topographic index (CTI) (or wetness index). It describes the effect of topography (slope and aspect) on the location and size of areas of water accumulation in soils. The wetness index is calculated by:

wT = ln (A / tan b)



wT:  Wetness index
A:   Specific catchment area (upslope area draining through a point(per unit contour length)
b:   Slope angle at the point

Hydrologically, the specific catchment area (A) is a measure of surface or shallow subsurface runoff at a given point on the landscape, and it integrates the effects of upslope contributing area and catchment convergence and divergence on runoff. The wetness index reflects the tendency of water to accumulate at any point in the catchment (in terms of A) and the tendency for gravitational forces to move that water downslope (expressed in terms of tan b as an appropriate hydraulic gradient). A GIS (geographic information system) can be used to calculate the wetness index based on DEM (digital elevation model) data to derive information about soil moisture. Furthermore, the wetness index can be used to derive the zones of saturation or variable source areas in a study area (catchment) and using a threshold wetness index saturation overland flow is derived . In the simulation model TOPMODEL, the CTI concept is combined with a storage model (storage for interception, infiltration, and the saturated zone) (Beven et al., 1984) . Care must be taken in applying a static wetness index to predict the distribution of a dynamic process like soil water content. A dynamic wetness index was developed by Barling et al. (1994) considering the change of soil moisture (i.e. wetness index) across time.

The depth of soil profiles also influences the soil moisture content. Thick soil profiles are able to store large amounts of water. The shallower a soil profile the less water can be stored . Such soils are prone to low soil moisture contents.

Soil texture influences the soil moisture content, where assumed the same climatic conditions, sandy soils have the tendency for low soil moisture content, silty soils for average soil moisture contents, and clayey soils for high soil moisture contents . This is due to the different pore space distribution in coarse and fine textured soils. Sandy soils have a high amount of macropores (pore diameter > 10 micrometer ), silty soils a high amount of medium-sized pores (pore diameter 0.2 – 10 micrometer), and clayey soils a high amount of fine pores (pore diameter < 0.2 micrometer) . The term soil moisture refers to the presence or absence either of ground water or of water held at a tension of less than 1500 kPa, in the soil or in a specific horizon, by periods of the year. Water held at a tension of 1500 kPa or more is not available to keep most plants alive. In Table 1. the soil moisture classes as defined in Soil Taxonomy are listed.

Table 1. A classification of soil moisture regimes.

Soil moisture regime Characteristics
Dry Soil moisture content less than the amount retained at 15 atmospheres of tension (1500 kPa – permanent wilting point). ‘In most years’ – 6 out of 10 years
Xeric Soils of temperate areas that experience moist winters and dry summers (i.e. mediterranean climates)
Aridic/Torric Soils are dry more than half the time (in arid climatic zone)
Perudic In most years precipitation exceeds evapotranspiration every month of the year
Udic In most years soils are not dry more than 90 consecutive days
Ustic In most years soils are dry for 90 consecutive days and moist in some part for half the days the soil temperature is above 5°C (i.e., during potential growing season)
Aquic Soils that are sufficiently saturated, reducing conditions occur. They usually have low chroma mottles or have gleyed subsoils

When soil moisture is high, as in wet or humid climates, there is a net downward movement of water in the soil for most of the year, which usually results in greater leaching of soluble materials, sometimes out of the soil entirely, and the translocation of clay particles from upper to lower horizons. In arid climates there is net upward movement of water in the soil, due to high evapotranspiration rates, which results in upward movement of soluble materials (e.g. salts). These accumulated materials can become cemented (-> pans), which are impenetrable to roots and lower infiltration tremendously.

Temperature varies with latitude and altitude, and the extent of absorption and reflection of solar radiation by the atmosphere. Solar radiation (direct radiation and diffuse radiation) increases with elevation, differs seasonally, and is influenced by cloud cover or other atmospheric disturbance (e.g. air pollution). The absorption of the solar radiation at the soil surface is affected by many variables such as soil color, vegetation cover, and aspect. In general, the darker the soil color, the more radiation is absorbed and the lower the albedo. The effect of vegetative cover on absorption varies with density, height, and color of the vegetation. Hence the absorption differs in areas with decidious trees (soil surface is shaded by trees most of the year) and arable land (soil surface is not shaded throughout the year). Light, or whitish-colored, soil surfaces tend to reflect more radiation. When incoming solar radiation is reflected, there is less net radiation to be absorbed and heat the soil. Snow is especially effective in reflecting the incoming solar radiation. Soil moisture controls also the heating up or cooling down of soils. Water has a high specific heat capacity (1 cal g -1 C), whereas dry soils have a specific heat capacity of about 0.2 cal g -1 C. This means that sandy soils cools and heats more rapidly than soils high in silt or clay. Once a wet soil is warmed, it takes longer to cool than a dry soil. In the Northern Hemisphere, south-facing slopes tend to be warmer and thus more droughty than north-facing slopes.

Temperature affects the rate of mineral weathering and synthesis, and the biological processes of growth and decomposition. Weathering is intensified by high temperatures, hence weathering is stronger in the tropics than in humid regions. Temperature also influences the degree of thawing and freezing (physical weathering) in cold regions. Biological processes are intensified by rising temperatures. Reaction rates are roughly doubled for each 10°C rise in temperature, although enzyme-catalyzed reactions are sensitive to high temperatures and usually attain a maximum between 30 and 35 °C.

From Dokuchaev on (about 1870), many pedologist in Europe and North America regarded climate as predominant in soil formation. The relationship between climatic zones and broad belts of similar soils that stretched roughly east-west across Russia inspired the zonal concept of soils. Zonal soils are those in which the climatic factor, acting on the soil for a sufficient length of time, is so strong as to override the influence of any other factor. Intrazonal soils are those in which some local anomaly of relief, parent material or vegetation is sufficiently strong to modify the influence of the regional climate. Azonal or immature soils have poorly differentiated profiles, either because of their youth or because some factor of the parent material or environment has arrested their development. In the U.S. the zonal concept was used in soil classification as published in the USDA Yearbook of Agriculture (Baldwin et al., 1938).

Table 2. Soil classification in 1938 USDA Yearbook of Agriculture (highest two categories only)

Category 6 – Order Category 5 – Suborder
Zonal soils Pedocals (soils with calcium carbonate accumulation)

Pedalfers (soils with iron and aluminum accumulation)

Soils of the cold zone: 1. Light-colored soils or arid regions 2. Dark-colored soils of the semiarid, arid, subhumid, and humid grasslands

3. Soils of forest-grassland transitions 4. Light -colored podzolized soils of the timbered regions 5. Lateritic soils of forested warm-temperate and tropical regions

Intrazonal soils 1. Halomorphic (saline and alkali) soils of imperfectly drained arid regions and littoral deposits 2. Hydromorphic soils of marshes, swamps, see areas, and flats 3. Calomorphic
Azonal No suborders

The concept of soil zonality is not very helpful when applied to soils of the subtropics and tropics. There, were land surfaces are generally much older than in Europe, and have consequently undergone many cycles of erosion and deposition associated with climatic change, the age of the soil and its topographical relation to other soils in the landscape are factors of major importance. The zonal concept is also of little use in regions such as Scandinavia or the Northern U.S., where much of the parent material of present-day soils is young (Pleistocene deposits) and relief plays a powerful role in soil formation.

Table 3. Definitions and features of soil temperature regimes.

Temperature regime Mean annual temperature in root zone [degree C] Characteristics and some locations
Pergelic < 0 Permafrost (i.e. the depth of freezing in winter exceed the depth of thawing in summer, as a consequence, a layer of permanently frozen soil of grounds develop) and ice edges common. Tundra of northern Alaska and Canada and high elevations of the Rocky Mountains.
Cryic 0 -8 Cool to cold soils of the Northern Great Plains of the U.S., forested regions of eastern Canada.
Frigid < 8 A soil with a frigid regime is warmer in summer than a soil with cryic regime. The difference between mean summer and mean winter soil temperatures is more that 5o C.
Mesic 8 – 15 Midwestern and Great Plains regions where corn and winter wheat are common crops.
Thermic 15 – 22 Coastal Plain of southeastern U.S. where temperatures are warm enough for cotton.
Hypothermic > 22 Citrus areas of Florida peninsula, southern California. Tropical climates and crops.

If the name of a soil temperature regime has the prefix iso, the mean summer and mean winter soil temperatures for June, July, and August and for December, January, and February differ less than 5 oC at a depth of 50 cm.

Temperature regime Mean annual soil temperature [degree C]
Isofrigid < 8
Isomesic >= 8 – 15
Isothermic >= 15 – 22
Isohyperthermic >= 22

Arnold J.G., Allen P.M., Bernhardt G. 1993. A Comprehensive Surface-Groundwater Flow Model. Journal of Hydrology, 142: 47-69.

Baldwin M., Kellogg C.E., and Thorp J., 1938. Soil Classification. Yearbook of Agriculture, U.S. Dept. Agric.,U..S.Govt. Printing Office, Washington, DC: 979-1001.

Barling R.D., Moore I.D., and Grayson R.B., 1994. A Quasi-dynamic Wetness Index for Characterizing the Spatial Distribution of Zones of Surface Saturation and Soil Water Content. Water Resources Research, 30 (4): 1029-1044.

Beven K.J., Kirkby M.J., 1979. A Physically Based, Variable Contributing Area Model of Basin Hydrology. Hydrological Sciences Bulletin, 24 (1): 43-69. Beven K.J., Kirkby M.J., Schofield N., Tagg A.F., 1984. Testing a Physically-Based Flood Forecasting Model (TOPMODEL) for Three U.K. Catchments. Journal of Hydrology, 69: 119-143.

Maidment D.R., 1993. Handbook of Hydrology, McGraw-Hill, Inc., New York.

Smith R.E., 1992. OPUS: An Integrated Simulation Model for Transport of Nonpoint-Source pollutants at the Field Scale. Vol I. Documentation, USDA-ARS-98.

Soil Conservation Service – U.S. Department of Agriculture. 1985. National Engineering Handbook, sec. 4, Hydrology, U.S. Government Printing Office, Washington D.C.

USDA-ARS. 1995. Water Erosion Prediction Project (WEPP). NSERL Report No. 11.


The soil and the organisms living on and in it comprise an ecosystem. The active components of the soil ecosystem are the vegetation, fauna, including microorganisms, and man.

Vegetation: The primary succession of plants that colonize a weathering rock culminates in the development of a climax community, the species composition of which depends on the climate and parent material, but which, in turn, has a profound influence on the soil that is formed. For example in the Mid-West of the U.S. deciduous forest seems to accelerate soil formation compared to grassland on the same parent material under similar climatic conditions. Differences in the chemical composition of leaf leachates can partly account for a divergent pattern of soil formation. For example acid litter of pines or heather favors the development of acid soils with poor soil structure, whereas litter of decidious trees favors the development of well structured soils.

Meso-/Macrofauna: Earthworms are the most important of the soil forming fauna in temperate regions, being supported to a variable extent by the small arthropods and the larger burrowing animals (rabbits, moles). Earthworms are also important in tropical soils, but in general the activities of termites, ants, and beetles are of greater significance, particularly in the subhumid to semiarid savanna of Africa and Asia. Earthworms build up a stone-free layer at the soil surface, as well as intimately mixing the litter with fine mineral particles they have ingested. The surface area of the organic matter that is accessible to microbial attack is then much greater. Types of the mesofauna comprise arthropods (e.g. mites, collembola) and annelids (e.g earthworms, enchytraeids).

Table 4. Earthworm biomass in soils under different land use (White, 1987)

Earthworm biomass [kg/ha]
Hardwood and mixed woodland: 370 – 680
Coniferous forest: 50 – 170
Pasture: 500 – 1500
Arable land: 16 – 760

Microorganisms: The organic matter of the soil is colonized by a variety of soil organisms, most importantly the microorganisms, which derive energy for growth from the oxidative decomposition of complex organic molecules. During decomposition, essential elements are converted form organic combination to simple inorganic forms (mineralization). Most of the microorganisms are concentrated in the top 15 – 25 cm of the soil because C substrates are more plentiful there. Estimates of microbial biomass C range from 500 to 2,000 kg /ha to 15-cm depth (White, 1987). Types of microorganisms comprise bacteria, actinomycetes, fungi, algae, protozoa, and soil enzymes.

Man: Man influences soil formation through his impact to the natural vegetation, i.e., his agricultural practices, urban and industrial development. Heavy machinery compacts soils and decreases the rate of water infiltration into the soil, thereby increasing surface runoff and erosion. Land use and site specific management (e.g. application of fertilizer, lime) also act on soil development.

White R.E., 1987. Introduction to the Principles and Practices of Soil Science. Blackwell Scientific Publ. , Palo Alto, CA.


Major topographical features are easily recognized in the field (e.g. mountains, valleys, ridges, crests, sinks, plateau, floodplains). For detailed description of topography Digital Elevation Models (DEMs) are used. In a DEM each pixel of a landscape is described by a data triplet consisting of Easting, Northing, and the elevation.

DEMs are available in different quality. Examples for DEMs are given in the following: U.S. Geological Survey (USGS) topographic 7.5 x 7.5-min blocks DEM (equivalent to a scale 1:24,000): Horizontal resolution = 30 m, root mean square error is generally =/- 7 m.

U.S. Geological Survey DEM: 1o x 1o blocks representing one-half of the standard 1:250,000 scale 1o x 2o quadrangle maps: Horizontal resolution = 90 m in the north-south direction and about 60 m in the east-west direction, accuracy for flat terrain +/- 15 m and for steep terrain 60 m.

High quality DEMs (e.g. horizontal resolutions of 5 – 10 m). An example for a high quality DEM is shown in Figure 2.

Figure 2. DEM for an experimental site at Arlington Agricultural Research Station in Southern Wisconsin.

Topographic attributes such as slope, aspect, specific catchment area, plan and profile curvature can be derived from DEMs using surface fitting functions of a GIS (geographic information system) or topographic programs such as TAPES-G (Gallant et al., 1996).

Table 5. Selected primary topographic attributes important in pedology.

Topographic attribute Definition Hydrologic significance
altitude elevation climate, vegetation type, potential energy
slope gradient overland and subsurface flow, velocity and runoff rate
aspect slope azimuth solar radiation
catchment area area draining to catchment outlet runoff volume
specific catchment area upslope area per unit width of contour runoff volume
flow path length maximum distance of water flow to a point in the catchment erosion rates, sediment yield
profile curvature describes the shape of a slope in a downward direction and indicates the rate of change in gradient water flow, flow velocity, sediment transport processes (erosion, deposition)
plan curvature describes the shape of the slope in a direction perpendicular to the slope and indicates the rate of change in gradient converging/diverging flow, soil water content

Figure 3. Landform elements of a hillslope.

Figure 4. Different shapes of a hillslope.

Based on topographic attributes landform elements can be delineated. Examples are given by Huggett (1975), Pennock et al. (1987), and Irvin (1996). They related landform elements to soil properties and hydrologic characteristics, which also influence soil genesis. Huggett combined vertical and horizontal slope curvatures (slope shapes) to predict soil drainage classes. He states that in general hydraulic conductivity decreases with depth. Thus, material and soil solution throughflow vary with both profile (downslope) and plan (across-slope) curvature. The water flux contains dissolved and suspended materials, which it moves from the upper reaches of the valley basin to lower parts. This movement may result in eluviation in the upper zone of the basin and illuviation along the lower reaches. Pennock et al. (1987) used combinations of gradient, plan and profile curvature to define distinct landform elements, which were related to soil moisture. The result of their study indicate that moisture content relates to elements in the sequence shoulder < backslopes < footslopes and divergent elements < convergent elements. Irvin (1996) related landform elements to soil properties (e.g. silt depth) in the Driftless Area of Southern Wisconsin. Generally, an increase in slope is associated with a reduction in:

  • Leaching
  • Organic matter content
  • Clay translocation
  • Mineral weathering
  • Horizon differentiation
  • Solum thickness

Topographic attributes and vegetation cover affect soil moisture by governing the proportions of surface runoff to infiltration. Soil with impermeable sub-soils and those developing on slopes, may show appreciable lateral subsurface flow. Thus, at the top of the slope, the soils tend to be freely drained with the water table at considerable depth, whereas the soils at the backslopes and footslopes are poorly drained, with the water table near or at the soil surface. The succession of soils forming under different drainage conditions on relatively uniform parent material comprises a hydrological sequence, an example of which is shown in Figure 5. As the drainability deteriorates, the oxidized soil profile, with its orange-red colors due to ferric oxides, is transformed into the mottled and gleyed profile of a reduced soil (soil color: gray, green).

The importance of relief was highlighted by Milne (1935) , who recognized a recurring sequence of soil forming on slopes in a generally undulating landscape. He introduced the term catena (Lat. ‘chain’) to describe a sequence of contiguous soils extending from hill to top of a hillslope.

Figure 5. Hydrological sequence of soils formed under major influence of relief: Soil profile #1 is well drained (summit), #2 moderately well drained (backslope), #3 poorly drained (footslope), and #4 very poorly drained (toeslope).

Each hillslope with a slope gradient is subdued to transport of soil particles. Erosion tends to be higher on convex sites with steep slopes compared to concave sites with low gradient. The soils at shoulders tend to be more shallow due to erosion, whereas the soils on footslope and toeslope areas tend to be thicker due to deposition. As erosion reduces the thickness of the A horizon, the upper part of the B horizon becomes incorporated into the lower part of the A horizon and the upper part of the C horizon becomes incorporated into the lower part of the B horizon. The sediment transport is different for each soil particle size. The transport of coarse-sized particles (sand) is lowest, whereas the transport of fine soil particles (clay) and medium-sized particles (silt) is higher. Clay particles form aggregates with organic material and iron and aluminum oxides, hence those aggregates are very stable and are less susceptible to sediment detachment. Coarse-sized particles are heavy and therefore also difficult to detach. Medium-sized particles (silt) are prone to erosion. If erosion occurs on a hillslope the silt content often is higher in the bottom soils compared to soils on the hillslope shoulder. Increasing the slope length allows water, which ran off the upper part of the slope to infiltrate in the lower part of the slope and to deposit eroded material carried in suspension.

Relief has also an important influence on the local climate and the vegetation. Changes in elevation affect the temperature (a decrease of approximately 0.5 degree C per 100 m increase in height), the amount and form of the precipitation and the intensity of storm events., thus affecting soil moisture relations. These factors interact to influence the type of vegetation.

Gallant J.C., and J.P. Wilson. 1996. TAPES-G: A Grid-based Terrain Analysis Program for the Environemental Sciences. Computers & Geosciences, 22 (7): 713-722.

Huggett R.J., 1975. Soil Landscape Systems: A Model of Soil Genesis. Geoderma 13: 1-22.

Irvin B.J. 1996: Spatial Information Tools for Delineating Landform Elements to Support Soil/Landscape Analysis. PhD Thesis, University of Wisconsin-Madison.

Milne G. , 1935. Some Suggested Units of Classification and Mapping, Particularly for East African Soils. Soil Research, 4: 183-198.

Pennock D.J., B.J. Zebarth, and E. de Jong. 1987. Landform Classification and Soil Distribution in Hummocky Terrain, Saskatchewan, Canada. Geoderma, 40: 297-315.

Parent Material

The nature of the parent material has a decisive effect on the properties of soils. Properties of the parent material that exert a profound influence on soil development include texture, mineralogical composition, and degree of stratification. Soil may form directly by the weathering of consolidated rock in situ (a residual soil), saprolite (weathered rock), or it may develop on superficial deposits, which may have been transported by ice, water, wind or gravity. These deposits originated ultimately from the denudation and geologic erosion of consolidated rock. Consolidated material is not strictly parent material, but serves as a source of parent material after some physical and /or chemical weathering has taken place. Soils may form also on organic sediments (peat, muck) or salts (evaporites). The chemical and mineralogical compositions of parent material determine the effectiveness of the weathering forces. During the early stages of soil formation, rock disintegration may limit the rate and depth of soil development. The downward movement of water is controlled largely by the texture of the parent material. Furthermore, parent material has a marked influence on the type of clay minerals in the soil profile.

Figure 6. Bedrock geology of Wisconsin.


Time acts on soil formation in two ways:

  • The value of a soil forming factor may change with time (e.g. climatic change, new parent material).
  • The extent of a pedogenetic reaction depends on the time for which it has operated.

Monogenetic soils are those that have formed under one set of factor values for a certain period of time. Soil that have formed under more than one set of factor values are called polygenetic.

Very old soils are formed on weathered consolidated rocks (e.g. granite, basalt), where the rocks were formed more than 500 million years ago (Paleozoikum). In Africa or Australia such old soils may be found.

The climate has changed over geological time, the most recent, large changes were associated with alternating glacial and interglacial periods of the Pleistocene. Europe and North America sustained four distinct ice invasions, whereas each glacial period was separated by long interglacial ice-free intervals. Those were times of warm or semitropical climate. The total length of the Pleistocene ice age is estimated 1 – 1.5 million years. The glaciers disappeared from Northern America approximately 12,000 years ago. As the glacial ice was pushed forward soil was swept away, hills were rounded, valleys filled and the underlying rocks were severely ground and gouged. Finally, when the ice melted a mantle of glacial drift remained, a new regolith and fresh parent material for soil formation. The influence of parent material is much more apparent in the soils of glaciated regions, where insufficient time has elapsed sine the ice retreated to permit the full development of soils.

One of our youngest soils are formed on alluvial or lacustrine materials generally have not had as much time to develop as the surrounding upland soils. Young in age are also colluvial soils, where sediment transport occurred recently.

In Figure 7 a hypothetical soil development across time is shown. The parent material might be relatively unweathered bedrock. After weathering of bedrock and the accumulation of organic matter at the soil surface there occurs the development of an A horizon, due to processes such as decomposition and mineralization. After an A horizon is formed slowly a B Horizon is developed, due to the formation of clay minerals (denoted by a lower case ‘t‘). In a humid environment such as the central United States material from the upper part of the soil profile is leached downwards (e.g. clays, organic material) and an eluvial horizon (E horizon) is formed. The accumulated material is precipitated in a horizon below the E , the so-called illuvial horizon, in this case the Bt horizon.

Figure 7. Stages of development of soils across time for a soil in the central United States under forest.

Primary Mineral Components of Soils

Primary Minerals

The thickness of the earth’s crust varies from 10 km under the ocean to 30 km under the continents. Of the 88 naturally occurring elements on earth, only 8 make of most of the crust. The earth’s crust and soils are dominated by the silicic acid in combination with Na, Al, K, Ca , Fe and O ions. In Table 1 the mean elemental content of soil and crustal rocks, and the soil enrichment factors are listed. Elements with high enrichment factors (EF) are C, N, S , and elements with low EF are Na, Mg , Al, P, Cl, K, Ca, Mn and Fe. The latter ones are important nutrients for plant growth.

Table 1. Mean elemental content of soil and crustal rocks, and the enrichment factors.

Element Mean elemental content of soil [mg/kg] Mean elemental content of crust [mg/kg] Enrichment Factor (EF)
Li 24 20 1.2
Be 0.92 2.6 0.35
B 33 10 3.3
C 25,000 480 52
N 2,000 25 80
O 490,000 474,000 1.0
F 950 430 2.2
Na 12,000 23,000 0.52
Mg 9,000 23,000 0.39
Al 72,000 82,000 0.88
Si 310,000 277,000 1.1
P 430 1,000 0.43
S 1,600 260 6.2
Cl 100 130 0.77
K 15,000 21,000 0.71
Ca 24,000 41,000 0.59
Sc 8.9 16 0.56
Ti 2,900 5,600 0.52
V 80 160 0.50
Cr 54 100 0.54
Mn 550 950 0.58
Fe 26,000 41,000 0.63
Co 9.1 20 0.46
Ni 19 80 0.24
Cu 25 50 0.50
Zn 60 75 0.80
Ga 17 18 0.94
Ge 1.2 1.8 0.67
As 7.2 1.5 4.8
Se 0.39 0.05 7.8
Br 0.85 0.37 2.3
Rb 67 90 0.74
Sr 240 370 0.65
Y 25 30 0.83
Zr 230 190 1.2
Nb 11 20 0.55
Mo 0.97 1.5 0.65
Ag 0.05 0.07 0.71
Cd 0.35 0.11 3.2
Sn 1.3 2.2 0.59
Sb 0.66 0.20 3.3
I 1.2 0.14 8.6
Cs 4.0 3.0 1.3
Ba 580 500 1.2
La 37 32 1.2
Hg 0.09 0.05 1.8
Pb 19 14 1.4
Nd 46 38 1.2
Th 9.4 12 0.78
U 2.7 2.4 1.1

Those elements are components of primary minerals, whereas primary minerals are components of parent rocks. There are almost 3000 known minerals, but only 20 are common and just 10 minerals make up 90 % of the earth’s crust. Primary minerals are defined as minerals found in soil but not formed in soil. This definition is different from that of secondary minerals , which are defined as minerals formed in soils.

Most of the primary minerals (primary silicates) have a crystalline structure, i.e., a structure in which ions are arranged in an orderly and repeated spatial pattern. The fundamental unit in silicates is the silicon-oxygen tetrahedron, which is composed of a central silicon ion surrounded by four closely-packed and equally-spaced oxygen ions. The four positive charges of Si 4+ are balanced by four negative charges from the four oxygen ions (O2 ), one from each ion, thus each discrete tetrahedron has four negative charges. The central ion may be either Al 3+ , Fe 2+, or Mg 2+. When in six-folded coordination, oxygens form an eight-sided octahedron. If larger Ca 2+, Na +, or K + ions are present, they occur at the center of clusters of tetrahedra, with each tetradedron supplying a part of all the oxygens needed for eight-fold or twelve fold coordination. In this arrangement, the larger cations provide a center of positive charge that attracts and holds the clusters of tetrahedra together. The cations occuring in this position, i.e., outside or between neighboring tetrahedra are called accessory cations. Si 4+ and Al3+ ions are small and have a high charge (valence). In general, the smaller the cation and the higher its valence the stronger the bond between it and the oxygen.

Figure 1. Silicon tetrahedron.
Figure 2. Aluminum tetrahedron.
Stability in minerals requires their structure to be electrically neutral, i.e., the negative charge of the O2- in the structure must be equally balanced by the positive charge of the cations. Isomorphous substitution is the replacement of an ion with higher valence by some other kind of cation. This process is supported by a high concentration of substituting ions in a mineral-forming medium so as to increase their chance of entering the mineral structure in place. The pattern of substitution is generally the following: Al3+ substitutes for Si4+, and Fe 2+ and Mg2+ substitutes for Al 3+. An electrical imbalance occurs because the valence of the substituting ions is lower than that of the ions in replace. Neutralization of the excess negative charge is accomplished by the inclusion of accessory cations in the structure. In the primary silicates, Ca 2+, Na+ , K + are the principal accessory cations that neutralize the negative charge resulting from ion substitution.

The most important primary silicates are discussed in the following:

Framework Silicates: They are composed of tetrahedra linked trough their corners into a continuous 3D-structure. Quartz is a framework silicate composed entirely of silicon-oxygen tetrahedra. The bulk density of quartz is 2.65 g/cm3 and quartz is highly resistant to mechanical abrasion and chemical weathering. Quartz is very common in most igneous, metamorphic and sedimentary rocks. In feldspars the Si4+ is partly replaced by Al 3+, which results in a positive charge balanced by Na +, K + or Ca 2+ ions. In the alkali feldspars Na+ and K +, and in the plagioclase, Na + and Ca2+ are the dominant accessory cations. Feldspars are the most abundant minerals in the earth’s crust; they make up 50 – 60 % of the crustal rocks.  

Figure 3. General structure of framework silicates.

Chain Silicates: The amphibole and pyroxene are chain silicates, whereas the Si4+ is also partly replaced by Al3+, but the chains are hold together by Na+, K +, Ca2+, Mg 2+, Fe2+, Al 3+ and /or Ti3+ ions. The tetrahedras are linked to each other by sharing two of the three basal corners to form continuous chains. The pyroxene are composed of silica tetrahedra which form single chains, whereas the amphiboles are composed of silica tetrahedra double chains. Because the bonding between chains is not strong the amphibole and pyroxene are easily weatherable.  

Figure 4. General structure of chain silicates.

Ortho- and Ring Silicates: They include the olivines, zircon, and titanite. In olivines the silicon-tetrahedra is arranged in sheets and linked by Mg2+ and / or Fe 2+ ions. Olivine is found in basalt and volcanic rocks.  

Figure 5. General structure of ring silicates.

Sheet Silicates: They are composed of three basic sheets: Silicon tetrahedral sheet: Composed of silicon-tetrahedra linked together in a hexagonal arrangement with the three basal oxygen ions of each tetrahedron in the same plane and all the apical oxygen ions in a second plane. Thus the silicon tetrahedral sheet is a hexagonal planar pattern of silicon-oxygen tetrahedra.

Aluminum hydroxide sheet: The basic unit of this sheet is the aluminum-hydroxyl octahedron in which each ion is surrounded by six closely packed hydroxyl groups, in such a way that there are two planes of hydroxyl ions, with a third plane containing aluminum ions sandwiched between the two hydroxyl planes. In order that all the valences of the structure be satisfied only two out of every three positions in the aluminum hydroxide sheet are occupied by aluminum ions forming what is known as a dioctahedral structure.

Magnesium hydroxide sheet: This has a similar structure to the aluminum hydroxide sheet but the aluminum is replaced by magnesium, and because magnesium is divalent all the sites in the middle plane are occupied, forming a trioctahedral structure.

In pyrophyllite one aluminum hydroxide sheet is lying between two silicon tetrahedral sheets and is known as a 2 : 1 type mineral. Micas are the most common primary sheet silicates, such as muscovite (white mica) and biotite (black mica). The micas contain oxygen in octahedra as well as in tetrahedra, with both occuring in a sheet like arrangement. Because of the ratio of two tetrahedral sheets to each octahedral sheet, the micas are called 2 : 1 layer minerals. Micas are generally found in granitic pegmatites, which are coarsely crystalline, igneous rocks. In muscovite one-quarter of the silicon ions is substituted by aluminum ions in the silicon tetrahedral layers. This imbalance in charges is satisfied by potassium, which bonds the composite sheets together. In biotite within the Mg-hydroxide sheet about one-third of Fe2+ is substituted for Mg2+. The negative charges occurring in this sheet are neutralized by potassium, which bonds the composite sheets together. The potassium is positioned in the interlayer space between neighboring layers. The potassium bonding is weak, where splitting may occur.

Figure 6. Diagrammatic structure of muscovite (mica). Table 2. Primary minerals.

Primary minerals Chemical formula Importance
Quartz SiO2 Abundant in sand and silt
Feldspar (Na,K)AlO2[SiO 2]3

CaAl2O4 [SiO2]2

Abundant in soil that is not leached
Mica K2Al2 O5[Si2 O5]3 Al 4(OH)4

K2Al2 O5[Si2 O5]3 (Mg,Fe) 6(OH)4

Source of K in most temperate zones
Amphibole (Ca,Na,K)2,3(Mg,Fe,Al) 5(OH)2[Si,Al4O11] 2 Easiliy weathered to clay minerals
Pyroxene (Ca,Mg,Fe,Ti,Al)(Si,Al)O3 Easily weathered
Olivine (Mg,Fe)2SiO 4 Easily weathered
Epidote Ca2(Al,Fe) 3(OH)SI3O 12 Highly resistant to chemical weathering, used as ‘index mineral’
Tourmaline NaMg3Al6 B3Si6 O27(OH,F)4 Highly resistant to chemical weathering
Zircon ZrSiO4 Highly resistant to chemical weathering
Rutile TiO2 Highly resistant to chemical weathering

Parent Rocks

The nature of parent material profoundly influences the characteristics of even highly weathered soils. Important for soil development is the chemical and mineralogical compositions of the parent material as well as the resistance of the material. Regolith is the loose, unconsolidated material (including soil) at the surface of the lithosphere. Most of the world’s soils have developed from sediments (transported unconsolidated material) that were originally derived from rocks such as glacial till, colluvium or loess. Soil development often occurs in a mixed heterogeneous material comprised of weathered bedrock and unconsolidated transported material. Saprolite is ancient residual soil and weathered rock formed by alteration of rock materials to clays and other residual material. A distinctive feature of saprolite is that the arrangement of the alteration products preserves the original rock structure and that the material has not been transported. For example, Saprolite is widespread on the Piedmont Plateau, where it has developed on various kinds of metamorphosed Paleozoic rocks (225 to 570 million years in age).

Other residual deposits comprise organic deposits, that are areas of marsh, swamp, and peat. Poorly drained areas in humid climates collect water and support vegetation adapted to wet environmental conditions (e.g. sedges). The term swamp usually applies to wet areas having trees. Either marshes or swamps may develop into peat bogs. Clayey or silty layers beneath such poorly drained ground are without oxygen, and that anaerobic condition weathers the mineral matter to sticky mud. Peat is the partly carbonized organic residue produced by decomposition of roots, trunks of trees, seeds, shrubs, grasses (reeds), ferns, mosses, and other vegetation. The decomposition is slowed because the ground is saturated with water, i.e., oxygen is excluded.

The mineral composition of parent rocks is responsible for the development of different soils. In general, the higher the content in calcium and magnesium and the lower the content of silicium in a parent rock the more likely soils with high base saturation are formed. Those soils are very productive, because of their high cation exchange capacity, which increase crop growth. Furthermore, calcium and magnesium form aggregates with clays, iron oxides and aluminum oxides, and organic matter. They improve soil structure, which is less prone to erosion.

Igneous Rocks Igneous rocks are formed by the solidification (hardening) of molten magma in the Earth’s crust. They vary in their composition of quartz and the light-colored Ca or K/Na silicates. Acidic rocks are those relatively rich in quartz and Ca or K/Na silicates with light colors (e.g. granite), basic rocks are those low in quartz but high in Ca or K/Na silicates (e.g. gabbro, basalt). Since the minerals in basalt weather more easily than those in granite, a finer-textured soil will develop from basalt. Major areas covered by igneous rocks are in Scandinavia, Canada and the Pacific coast of North America. The complete weathering of minerals in basalt in humid tropical regions produces soils composed only of clay-sized particles. Based on the mineral composition of the rocks soils formed on granite develop less fertile soils, whereas soils formed on basalt develop soils with more bases and therefore a higher fertility. Table 3. Mineral composition of some igneous rocks [%] (Scheffer et al., 1989).

Granite Gabbro Basalt
SiO2 73.9 48.4 50.8
TiO2 0.2 1.3 2.0
Al2O3 13.8 16.8 14.1
Fe2O3 0.78 2.6 2.9
FeO 1.1 7.9 9.0
MnO 0.05 0.18 0.18
MgO 0.26 8.1 6.3
CaO 0.72 11.1 10.4
Na2O 3.5 2.3 2.2
K2O 5.1 0.56 0.82
H2O+ 0.47 0.64 0.91
P2O5 0.14 0.24 0.23

Sedimentary Rocks
These are composed of the weathering products of igneous and metamorphic rocks and are formed after deposition by wind and/or water. Cycles of geologic uplift, weathering, erosion and subsequent deposition of eroded materials in rivers, lakes and seas have produced superimposed strata of sediments. Under the weight of the overlying sediments, deposits gradually consolidate (diagenesis) to form sedimentary rocks. Often calcium (Ca) and magnesium (Mg) cement sedimentary rocks. A characteristic of sedimentary rocks is their stratification. Sedimentary rocks cover approximately 75 % of the earth’s surface.

There are three major groups of sedimentary rocks:

  1. Clastic Sedimentary Rocks
    They are composed of fragments of the more resistant minerals. Dependent on the size of the deposited material (conglomerates, sand, silt, clay) different clastic sedimentary rocks are developed. The percentage of quartz is higher than 75 % in sandstones and they weather to produce sandy soils. Graywackes are high in mica and chlorite and quarzites are high in silicium (Si). Sedimentary rocks with a high fraction of clay were originated by deposition of clays in slow running water. The clays are shaped like flat leaflets and therefore under pressure they become stratified to form clay slate (shale). Soils formed on clay slate do have a high content of bases, but they are less permeable due to the high clay content and difficult to use in agricultural production. Clay slate is not very resistant to weathering.  
  2. Chemical Sedimentary Rocks
    They are formed by precipitation or flocculation from solution, most commonly limestone and chalks. Limestones are high in calcium, whereas dolomites are high in calcium and magnesium. Soils developed from limestone are usually fine-textured soils.  
  3. Biogenic Sedimentary Rocks
    They develop when organisms living in oceans or lakes die, sink down and become consolidated. For instance, shell limestone and marl with mollusk shells are biogenic sedimentary rock high in calcium.

Metamorphic Rocks
Igneous rocks and sedimentary rocks that are subjected to intense heat and great pressure are transformed into metamorphic rocks. Changes in mineralogy and rock structure generally render the metamorphosed rock more resistant to weathering. Examples for metamorphic rocks are marble, schist, and gneis. Marble, the metamorphic equivalent of limestone or dolomite, maybe either coarse- or fine-grained. Gneiss, which is metamorphosed igneous or sedimentary rock, consists of layers of coarse grains, usually quartz or feldspars, alternating with layers of fine-grained minerals. Schist, metamorphosed shale, consists of micaceous layers. Metamorphic rocks are common in Africa, Brasil, western Australia and India.

Scheffer F., and Schachtschabel P. , 1989. Lehrbuch der Bodenkunde, Ferdinand Enke Verlag.

Transported Parent

Material Transported parent material is unconsolidated material (the regolith) that has undergone transport processes by wind, water or gravity.

Eolian (Wind)
Wind moves rock fragments by processes of rolling, saltation and eolian dispersion. The transport is different for clay, silt , and sand, dependent on the size of particles moved. For example, wind deposits are the sand dunes of shorelines and desert areas. Material swept away from dry periglacial regions has formed eolian deposits called loess in North America, central Europe, and China. It covers eastern Nebraska and Kansas, southern Wisconsin, southern and central Iowa and Illinois, northern Missouri, and parts of southern Ohio and Indiana, besides a wide band extending southward along the eastern border of the Mississippi River. During the glaciation, much fine material was carried miles below the ice sheets by streams. This sediment was deposited over wide areas by the overloaded rivers. Those areas without vegetation were prone to wind erosion when the climate became drier. This blown material is called loess. Loess is usually silty in character and has a yellowish color. Quartz seems to predominate, but large quantities of feldspar, mica, and pyroxene are found, too. Most loess is highly calcareous. Soils formed in loess are highly productive soils. Soils formed in loess material are susceptible to erosion, which might be a problem in hilly areas. Especially, in the Midwest where a combination of humid climate, prairie vegetation and loess material formed one of the best soils in the world.    

Figure 7. Approximate distribution of loess in the United States (after Brady, 1984)

Figure 8. Loess depositions in the vine growing area Kaiserstuhl, Southern Germany.  

Alluvial (Water, River)
Transport by water produced alluvial, terrace and footslope deposits. During the transport the rock material is sorted according to size and density and abraded, so that fluviatile deposits characteristically have smooth, round pebbles. The larger particles are moved by rolling or are lifted by the trubulence of the water. The smallest particles are carried in suspension. Colloidal material may not be deposited until the stream discharges into the sea, when flocculation occurs and estuarine deposits form.   Alluvial deposits are scattered along the borders of streams during flooding, whereas the material is stratified. When a flooding stream overflows its banks, coarse particles are settled close to the stream, whereas finer particles are deposited further away from the stream. Terraces are developed from floodplains as streams cut deeper channels. Several terraces may be found along a stream. Streams flowing from hills or mountains into dry valleys or basins drop their sediments in a fan-like deposit as the water spreads out. These alluvial fans are usually well drained and coarsely textured, being composed of sands and gravels. Sediments not deposited as floodplains are carried to the lake, gulf, or other bodies of water into which a stream empties. The decrease in velocity at the stream’s mouth results in the deposition of much of the suspended material, thus producing deltas, which are poorly drained. Floodplains as well as deltas are generally rich in plant nutrients and comparatively high in organic matter content. Terraces and alluvial fans, on the other hand, are more likely to be less fertile.  

Figure 9. Alluvial deposits: Deposition of fine and coarse sized particles if flooding occurs.

Lacustrine (Water, Lake)
When particles settle down in lakes lacustrine deposits occur. Mixed with organic material (e.g. skeleton of organisms) they may become cemented and form biogenic sedimentary rock. Glaciolacustrine deposits are the most common. Other kinds of lake deposits are small. One common kind is found in karst terrain, where limestone roofs over cavernous limestone have collapsed (e.g. in Florida). Still other kinds of lakes develop cut-off meanders at river beds. Crater lakes are formed in volcanic craters or calderas.

Marine (Water, Ocean)
It is common to find marine deposits along coastlines. This material was derived from sediments carried by streams and deposited in the ocean and gulf through decreased current velocity. Also, considerable debris is torn from the shoreline by the pounding of the waves and the undertow of the tides. If there have been changes in shoreline, the alternation of beds will show no regular sequence and considerable variation in topography, depth, and texture. These deposits have been extensively raised above sea level along the Atlantic and Gulf coast of the United States. Marine sediments are generally sandy and low in mineral nutrients.

Glacial (Ice)
During the Pleistocene (1.5 – 10,000 B.C), northern America and northern and central Europe, and parts of northern Asia were invaded by a succession of great ice sheets. As the glacial ice pushed forward, it conformed to the unevenness of the areas invaded. The mantle of soil was swept away and the underlying rocks were severely ground and gouged. Thus, the glacier became filled with rock fragments, carrying much of its surface and pushing great masses ahead. Finally, when the ice melted and the region was free, a mantle of glacial drift remained, a new regolith and fresh parent material for soil formation. The area covered by glaciers in North America is estimated at 10.4 million km2 and about 20 % of the U.S. is influenced by the deposits.   Glacial till is the material deposited directly by the ice. It is a mixture of rock debris of great diversity, especially the particle size range is very large. Glacial till is found mostly as deposits called moraines composed of unassorted material. The terminal moraines characterize the southernmost extension of the various glacial lobes. The ground moraine, a thinner and more level deposition laid down as the ice front retreated rapidly. It has the widest extent of all glacial deposits. An outstanding feature of glacial till material is their variability. This is because of the diverse ways by which the debris was laid down, of differences in the chemical composition of the original materials, and of fluctuations in the grinding action of the ice. The soils formed in such material are most heterogeneous.   The outwash plains are formed by streams heavily laden with glacial sediments. This sediment is usually assorted. Such deposits are particularly important in valleys and on plains, where the glacial waters were able to flow away freely. In many cases the ice front came to a standstill where there was no such ready escape for the water, and ponding occurred as a result of damming action of the ice. Often very large lakes were formed that existed for many years. The deposits that were made in these glacial lakes range from coarse materials near the shore to fine silts and clay in the deeper and stiller waters (glaciolacustrine deposits). As a consequence, the soils developed from these lake sediments are most heterogeneous.   It is customary to designate all the material deposited by glaciers and their melt water as glacial drift.

Figure 10. Glacial deposits: Glacial till, outwash, and loess.

Figure 11. Areas in the U.S. covered by the continental ice sheet and the deposits either directly from, or associated with, the glacial ice (after Brady, 1984).  

Mass movement
Erosion by water may be divided into four categories: (1) splash, (2) sheet, (3) rill, and (4) gully erosion. Sheet erosion refers to the uniform removal of soil from the surface, whereas rill erosion occurs concentrated in rills (or gullys), i.e., unevenly distributed across an area. The deposited material is called colluvium. Colluvium includes only those deposits that are or have been moving slowly downslope by ground creep. It is generally a chaotic mixture of coarse and fine-grained materials. Thickness of colluvium are generally less than 3 m and rarely more than 8m. Solifluction is a process which occurs on partly frozen soils, where the lower part of the soil profile is frozen and the upper part is highly saturated by water. The upper part of the profile flows slowly downhills. Mudflows (fine-grained debris) and avalanche (coarse-grained debris) are extreme events of mass movement acting upon gravity force.  

Figure 12. Colluvium deposits.  

Volcanic Activity
Volcanic ash is an amorphous (non-crystalline), fine, dustlike material thrown out of volcanoes. Ash falls on the surrounding land to form thick sediments for soil development. Other materials resulting from volcanic events are scoria, pumice, and bomb.

Significance in Pedology
The symbol ‘w‘ is used for horizons with development of color and structure. For example, a ‘Bw’ denotes a B horizon that has developed structure or color different, usually redder, than that of the A and C horizons but do not have apparent illuvial accumulations. Buried horizons which might develop due to erosion and deposition are designed by a ‘b‘. A ‘ r‘ denotes weathered bedrock. This symbol is only used with a master C horizon. It designates saprolite or dense till that is hard enough that roots only penetrate along cracks, but which is soft enough that it can be dug with a spade or shovel.   

Brady N.C., 1984. The Nature and Properties of Soils. MacMillan Publishing Company. Inc., New York.


Physical Weathering

Weathering in general refers to a group of processes by which surface rock disintegrates into smaller particles or dissolve into water due to the impact of the atmosphere and hydrosphere. The weathering processes often are slow (hundred to thousands of years). The amount of time that rocks and minerals have been exposed at the earth’s surface will influence the degree to which they have weathered.

Weathering processes are divided into three categories:

  • physical weathering
  • chemical weathering
  • biological weathering

Primary minerals and rocks are disintegrated in fragments due to physical weathering. This leads to environmental conditions (e.g. a higher surface area) that favor chemical weathering. There are several forms of physical weathering:

Abrasion: Water carrying suspended rock fragments has a scouring action on surfaces. Examples are the grinding action of glaciers, gravel, pebbles and boulders moved along and constantly abraded by fast-flowing streams. Particles carried by wind also have a ‘sand-blasting effect’.

Wetting and drying: Water penetrates into rocks and reacts with their constituent minerals.

Freezing and thawing: When water is trapped in the rock (or in cracks) repeatedly freezing and thawing results in forces of expansion and contraction (when water freezes, the increase in its volume is about 9 %).

Thermal expansion and contraction of minerals: Rocks are composed of different kind of minerals. When heated up by solar radiation each different mineral will expand and contract a different amount at a different rate with surface-temperature fluctuations. With time, the stresses produced are sufficient to weaken the bonds along grain boundaries, and thus flaking of fragments.

For instance, the difference in temperature in desert environments or mountain regions may range from 30 – 50 degrees C between day and night. Rocks are heated and cooled from the outside by change in solar radiation, which results in high temperature gradients inside and outside of the rocks (the heat conductivity of rocks is very low).

Pressure unloading or pressure-release jointing: There is a reduction in pressure on a rock due to removal of overlying material. This allows rocks to split along planes of weakness, called joints.

Crystallization: In arid environments, water evaporates at the surface of rocks and crystals form from dissolved minerals. Over time, the crystals grow (They expand their volume) and exert a force great enough to separate mineral grains and break up rocks.

Action of organisms: They aid in the physical disintegration of rocks.

Plant roots: They aid in the physical disintegration of rocks. Pressures exerted by roots during growth are able to rupture rocks.

Chemical Weathering

The difference between physical and chemical weathering is that with the latter one the mineral composition of the mineral or rock is changed. The larger the surface area, i.e, the smaller the fragments, the better for chemical weathering. Water is the dominant agent because it initiates chemical weathering. In the following there is a brief description of the most important chemical weathering processes:

Hydration: Ions have the tendency to hydrate when H 2O is present and dissociate. This kind of weathering happens in arid environments where salts are present. For example, chlorides and sulfates weather due to hydration. In general, ions with the same charge but smaller ion radius have a larger layer of H2 O ions and therefore do not tend to adsorb tight. The small Li + ion tends to remain hydrated at the surface, whereas the large Al3+ ion tends to dehydrate and become tightly adsorbed. The strength of adsorption increases in the following sequence:

Li+x x Na+ x x K+ x x Mg2+ x x Ca2+ x xAl3+

Hydrolysis: Water molecules at the mineral surface dissociate into H+ and OH and the mobile H+ ions (actually H 3O+) penetrate the crystal lattice, creating a charge imbalance, which causes cations such as Ca 2+, Mg2+, K + and Na+ to diffuse out. For example, the feldspar orthoclase hydrolyses to produce a weak acid (silicic acid), a strong base (KOH), and leaves a residue of clay mineral illite, which is a secondary mineral:

3KAl4 + Si3 O8 + 14H2 O <- -> K(AlSi3) 4Al2 4 O10(OH) 2 + 6Si(OH)4 + 2KOH

In hydrolysis reactions it has to be taken into account the important role played by dissolved CO2. This is shown in the hydrolysis of Mg-olivine:

Mg2SiO4 + 4CO2 + 4H2 O <- -> 2Mg2+ + 4HCO 3 + H 4SiO4

This reaction uses an acid (carbonic acid – H2 CO3)and therefore the solution becomes increasingly alkaline during completion of hydrolysis reactions.

Oxidation-Reduction: Several primary minerals contain Fe 2+ and Mn2+. If there are oxidizing environmental conditions the Fe2+ is oxidized to Fe3+ (precipitates as an insoluble oxyhydroxide, usually either ferrihydrite or the stable mineral goethite) and Mn2+ to Mn 3+ or Mn4+ partly inside the minerals, which results in a positive charge and the mineral becomes unstable. This charge imbalance is neutralized by a loss of some oxidized iron and manganese ions and/or some cations dissociate from the mineral. The precipitate may form a coating over the mineral surface, which slows down the subsequent rate of hydrolysis. Note that the oxidation of Fe 2+ to Fe3+ according to:

Fe2+ + 2H2 O + 1/2O2 < - -> Fe(OH) 3 + H+

is an acidifying reaction (acid solution weathering). The H + ions produced by this reaction will generally accelerate the rate of hydrolysis.

Complexation: Metals released from primary minerals such as Fe, Mn, and Al, build complexes with organic components, such as fulvic acids and humic acids, which are very stable. Important referring to chemical weathering is the loss of the cations out of the active system, therefore causing an imbalance between cations and anions.

Figure 1. Chemical weathering processes.


Weathering of primary minerals produce secondary minerals. Elements released from primary minerals are prone to leaching if they do not form complexes. The area of weathering is depleted first by Na+ , Ca2+, and Mg2+.

Biological Weathering

Lichens play an important part in weathering, because they are rich in chelating agents, which trap the elements of the decomposing rock in organo-metallic complexes. Some of the lichens being epilithic (i.e. living on the rock surface), some endolithic (actively boring into the rock surface), and some chasmolithic (living in hollows or fissures within the rock). Evidence for the operation of these processes comes mainly from detailed microscopic and microchemical analyses of the lichen : rock interface. The mechanisms and results of their actions is summarized in Table 1.

Table 1. Lichen weathering mechanisms and forms produced (Robinson et al., 1994).

Mechanisms Results of weathering
Chemical mechanisms

  • Chelation of extracellular, soluble compounds
  • Attack by oxalic acid
  • Attack by water acidified by respired carbon dioxide
Results of chemical action

  • Grooves at endolithic thalli interfaces
  • Etching minerals
  • Precipitation of alteration products, e.g. calcium oxalate, which may or may not play a further role in weathering
Physical mechanisms

  • Rhizine penetration
  • Thallus expansion and contraction on wetting and drying
Results of physical action

  • Exfoliation of rock surface layer
  • Cracking of rock
  • Increase in pore volume

Weathering Resistance

The resistance to weathering, i.e. the mineral stability of parent material depends on:

  • Types of mineral present
  • Surface area of rock exposed
  • Porosity of rocks

Weathering is not only dependent on the mineral composition but also on the porosity of the rock. Rocks consisting of coarse fragments (e.g. granite) easily weather physically but do not weather chemically fast. In contrast, rocks consisting of fine fragments (e.g. basalt) chemical weathering is higher than physical weathering. The weathering of stratified sedimentary rocks is dependent on the orientation of the stratification and the cementation.

In general, the resistance of a primary mineral to weathering increases with the degree of sharing of oxygens between adjacent Si tetrahedra in the crystal lattice. The SI-O bond has the highest energy of formation, followed by the Al-O bond, and the even weaker bonds formed between O and the metal cations (e.g. Na+, Ca 2+). In Figure 2 the ranking of some primary minerals in order of increasing stability is shown. Olivine weathers rapidly because the silicon tetrahedra are only held together by O-metal cations. In contrast quartz is very resistant because it consits entirely of linked silicon tetrahedra. In the chain (amphiboles and pyroxenes) and sheet (phyllosilicates) structures, the weakest points are the O-metal cation structures. Isomosphous substitution of Al3+ for Si4+ also contributes to instability because the proportion of Al-O to Si-O bonds increases and more O-metal cations bonds are necessary. This accounts for the decrease in stability of the calcium feldspars when compared with the sodium and potassium feldspars.

weak stability xOlivine, Ca2+ -Plagioclasex



xNa+-Plagioclase x

xK+-Plagioclase x

xMica (Muscovite)x

xQuartzx high stability
Figure 2. Stability of some primary minerals.

The rate of weathering is influenced by:

  • Temperature
  • Rate of water percolation
  • Oxidation status of the weathering zone

Weathering is dependent on the climate, i.e., the temperature and the mean annual precipitation rates resulting in different soil moisture contents. The mean lifetime of one millimeter of different rocks into a kaolinitic saprolite is shown in Table 2. These numbers exhibit that in cold, temperate, or tropical humid zones, the climate (temperature and precipitation) controls the rate of weathering.

Table 2. Mean lifetime of one millimeter of fresh rock (Nahon, 1991).

Rock Type Climate Lifetime (years)
Acid rocks tropical semi-arid

tropical humid

temperate humid

cold humid

65 to 200

20 to 70

41 to 250


Metamorphic rocks temperate humid 33
Basic rocks temperate humid

tropical humid



The oxidation status influences the degree of chemical weathering processes. An oxidizing environment favors the oxidation of ions such as Fe 2+ and Mn2+. Water is the agent forcing the processes of hydration and hydrolysis. High water contents mean also reducing (anaerobic) environmental conditions, which decrease the rate of oxidation.

Nahon D.B., 1991. Introduction to the Petrology of Soils and Chemical Weathering. John Wiley & Sons, Inc., New York.

Secondary Silicates


The structure of clay silicates is similar to that of primary silicates, i.e., they are sheet silicates. Secondary minerals are composed of silicon tetrahedral sheets, aluminum hydroxide sheets, and / or magnesium hydroxide sheets. Under mild (generally physical) weathering conditions, secondary minerals may be inherited as colloidal fragments of primary layer silicates, such as the micas. Under more intense weathering, the primary minerals may be transformed to secondary clay minerals, as when soil hydrous micas and vermiculites are formed by the leaching of interlayer K from primary micas. Neoformation of clay minerals is a feature of intense weathering, when minerals completely different from the original primary minerals are formed. The term ‘clay’ is used also to identify mineral particles of the size < 0.002 mm. Following, the most important secondary minerals are presented briefly: Kaolinites:They are composed of one tetrahedral sheet linked to an octahedral sheet, therefore they are classified as 1 : 1 type layer silicates. The two surfaces of a 1 : 1 mineral are formed by different ions: One consists of tetrahedral oxygens and the other of OHions belonging to the octahedral sheet. When the 1 : 1 sheets occur in stacks, the OH ions of one sheet lie next to and in close contact with the O2-layer of its neighbor. Because of this arrangement, the positive charge of the H+ ions in the OH layer exerts a strong attraction for the negative oxygens of the neighboring sheets. In this way the platelets of kaolinite are tightly bound together. Kaolinite is a non-expanding mineral, hence it is unable to absorb water into the interlayer position. The non-expanding character of kaolinite explains the failure of soils high in this clay to swell or shrink much on wetting and drying. Kaolinite has a basal spacing fixed at 0.72 nm, which is small compared to the other clay minerals.


General structure Diagrammatic representation

Montmorillonites (Smectite Group): These clay silicates form by crystallization from solution high in soluble silica and magnesium. Montmorillonite has a 2 : 1 layer structure. All tetrahedra in the sheets contain Si 4+ ions. Aluminium is the normal ion in the central sheet, but about one-eight of the octahedra contain Mg2+ as a substituting ion for Al3+. The negative charge caused by substitution is neutralized by various hydrated cations adsorbed to the surface of the sheets. The force of bonding between cations and the sheets is not very strong and depends on the amount of water present. In dry montmorillonites the bonding force is relatively strong. When wet conditions occur, water is drawn into the interlayer space between sheets and causes the clay to swell dramatically (expanding clay). A characteristic feature of montmorillonite is the extensive surface for the adsorption of water and ions, therefore the cation exchange capacity of montmorillonite is very high. Layers of the smectite group range in thickness from 0.98 to 1.8 nm or more.


General structure Diagrammatic representation

Vermiculites: These clays have a 2 : 1 structure of primary mica minerals. Vermiculites contain either Al3+ or Mg 2+ and Fe2+ as normal octahedral ions, and tetrahedral sheets in which Al 3+ occurs as a substituted ion in place of some of the Si 4+. Vermiculite differs from the micas in that it contains hydrated cations rather than unhydrated K+ in the interlayer space. The weak bonding afforded by these ions allows vermiculite to expand on wetting. Expansion is less than in montmorillonite, however. Unlike montmorillonite and kaolonite, vermiculite does not form by crystallization from solution, but instead it is formed by alteration, or the selective replacement of ions in a structure without destroying the structure (e.g. micas are altered to vermiculites). Layer spacing ranges from 1.0 to 1.5 nm or more.


General structure Diagrammatic representation

Hydrous Micas (Illites): They are 2 : 1 type minerals containing sufficient interlayer K+ to limit expansion on wetting. The K+ content of hydrous mica is less than that of micas. Charges not neutralized by K + are countered by hydrated cations. Formation of hydrous mica is favored in K-rich sediments. The process of hydrous mica formation is initiated as K+ replaces some of the interlayer cations of montmorillonites or vermiculites, and is completed when heat and pressure cause the dehydration and collapse of the clays into non-expanded forms. Hydrous micas are widespread in soils. The layer thickness of hydrous micas are about 1.0 nm.

Hydrous Mica:

General structure Diagrammatic representation

Chlorites: This group embraces a range of minerals that have certain outstanding characteristics in common. All have a basic 2 : 1 layer structure, and they are non-expanding. Chlorites differ from other 2 : 1 layer minerals in one unique respect, i.e., they contain a stable, positively charged octahedral sheet rather than adsorbed cations in the interlayer space. The octahedral sheet consists of two layers of OHions that enclose either Mg2+, Fe 2+, or Al3+ as the central octahedral cations and leads to a positive charge of the sheet. By virtue of its positive charge, the interlayer sheet neutralize the negative charge of the 2 : 1 sheets. Because chlorite contains two octahedral sheets, it is called a 2 : 1 : 1 layer mineral. Sometimes, octahedral materials in chlorite neither totally fill the interlayer space between sheets nor completely neutralize the negative charge of the sheets. This unsatisfied charge is neutralized by various cations adsorbed to the particle surfaces from the solution phase. The thickness of the chlorite layer is 1.4 nm.


General structure Diagrammatic representation

Allophanes: These are poorly categorized substances, which are sometimes regarded as clay minerals and at other times considered among hydroxides. It is for sure that they are poorly structured, i.e., amorphous in character and they consist of silica and hydrous oxides. They are abundant in soils derived from volcanic ash deposits. The structural formula of Allophane can be written as Si3Al 4O12*nH 2O. Another similar silicate is Imogolite which is Si 2Al4O 10*5H2O.

It should be stressed that clay minerals scarcely occur in pure form in soils, but they are a mixture of the clays presented in theoretical form above.


The properties of clay minerals are summarized in Table 1 and discussed in the following.

Table 1. Summary of clay mineral properties.

Secondary mineral Type Interlayer condition / Bonding CEC [cmol/kg] Swelling potential Specific surface area [m2/g] Basal spacing [nm]
Kaolinite 1 : 1 (non-expanding) lack of interlayer surface, strong bonding 3 – 15 almost none 5 – 20 0.72
Montmorillonite 2 : 1 (expanding) very weak bonding, great expansion 80 – 150 high 700 – 800 0.98 – 1.8 +
Vermiculite 2 : 1 (expanding) weak bonding, great expansion 100 -150 high 500 – 700 1.0 – 1.5 +
Hydrous Mica 2 : 1 (non-expanding) partial loss of K, strong bonding 10 – 40 low 50 – 200 1.0
Chlorite 2 : 1 : 1 (non-expanding) moderate to strong bonding, non-expanding 10 – 40 none 1.4
Allophane 10 – 50

All clay minerals show different expansions, whereas kaolinite, hydrous mica, and chlorite are non-expanding minerals and the others are expanding minerals. In kaolinite the bonding is strong because of tight H-OH bonding between the layers. The interlayer bonding of hydrous mica is mostly by K+ ions which is relatively strong. Montmorillonite and vermiculite show very weak to weak bonding due to various cations between the sheets, therefore they show a great expansion, especially in wet conditions. In chlorite the bonding is moderate to strong because of the positively charged octahedral layer.

The smaller the size of a fragment, the greater the ratio of its surface to volume, which defines the specific surface area. The specific surface area is low for kaolinite and hydrous mica and high for montmorillonite, vermiculite, and allophanes.This is because the surface area outside of the silicates (external surface) is increased by the surface area between the sheets, called interlayer area (or internal surface). In comparison coarse sand has a specific surface area of about 0.01 m2 /g, fine sand 0.1 m2/g, silt 0.1 – 1 m 2/g, and humic acids 800 – 1000 m2 /g (White, 1987).

The cation exchange capacity (CEC) is quite variable within and between mineral groups. For example, the CEC of the kaolinites, hydrous micas, chlorides, and allophanes is relatively low, whereas the CEC is high for montmorillonites and vermiculites. For comparison humic acids show the highest CEC in soils with 180 – 300 cmol/kg. There are two mechanisms, which are driving the CEC, both associated with the negative charges of silicate clays. The first involves unsatisfied valences at the broken edges of the silica and alumina sheets. Also, the flat external surfaces of the minerals have some exposed oxygen and hydroxyl groups, which act as negatively charged sites. Especially at high pH, the hydrogen of these hydroxyl dissociates slightly and the colloidal surface is left with a negative charge carried by the oxygen. In moderately to strongly acid soils the hydrogen is apparently tightly held and not subject to ready replacement by other cations. The magnitude of this pH-dependent charge varies with the type of the colloid. It accounts for most of the charge of the 1 : 1 type minerals and up to one fourth of that of some 2 : 1 type minerals. The cation exchange capacity of soils with a high amount of 2 : 1 type clays originates mainly from isomorphous substitution. These negatively charged sites are not affected by pH and constitute the permanent charge. Minerals such as kaolinite with their absence of isomorphous replacement have their exchange sites confined to the broken edges of crystals, therefore the CEC is low. On the other hand montmorillonite and vermiculite have a relatively large amount of isomorphous replacement resulting a large number of exchange sites and a high CEC.

Flocculation and dispersion are further important characteristics of clay minerals. Flocculation is the process, where the individual particles of clay are coagulated to form floccular aggregates. The degree and permanence of flocculation depend upon the nature of the ions present. For example, calcium and hydrogen tend to increase flocculation. Dispersion is defined as a process in which the individual particles are kept separate from one another. This is accomplished by potassium and sodium. Thus, depending upon the cations present in a soil, it may be either in a flocculated (aggregated) or in a dispersed (massive) state. Sodium saturated clays have a thick electric double layer surrounding the ion, that means the clays remain is suspension. Calcium suppresses the double layer and cause flocculation, while tri- and tetravalent ions are more efficient in causing flocculation. Clay translocation is closely related to flocculation and dispersion, respectively. The movement of clay requires that the clay be dispersed so that it can remain in suspension and be transported by water moving through pores and cracks in soil.


The usual source of kaolinites and montmorillonites is the precipitation at weathering sites. Hydrous micas are formed due to alteration of vermiculite or montmorillonite, so is the formation of vermiculite (alteration of mica or hydrous mica). Chlorites are formed by alteration of vermiculite or montmorillonites or in metamorphic rocks.

The stages in weathering are listed in Table 2. It is apparent that the composition of weathering solutions is strongly dependent on minerals that are undergoing weathering. First the original minerals dissolve and secondary minerals can form from it. Leaching of elements such as calcium, magnesium, sodium, potassium, and soluble silica supports further transformation processes. The gradual loss of soluble silica results in the formation and disappearance of clays in an ordered sequence, starting with those highest in silica content and ending with those containing no silica, i.e., the hydrous oxides. Over long periods the clays that form first eventually become unstable, decompose and they are replaced by other secondary clay minerals which are more stable. First 2 : 1 clay minerals are formed. Iron oxides may also appear early and they seem to persist almost infinitely in the weathering environment, which attests to their great stability under most conditions. As weathering proceeds kaolinites appear, or even kaolinites are decomposed, the silica released from it is leached, and the aluminum transforms to a hydrous oxide, usually gibbsite. These minerals tend to persist as the final products of long and intense silicate mineral weathering. The stages of weathering are time related functions, whereas the rate of weathering depends primarily on the climatic factors (temperature, precipitation). Silicate mineral weathering and clay synthesis are limited under either dry or cold conditions, but they proceed rapidly under hot, wet conditions, as in tropical regions. The time span required for a full weathering cycle shown in Table 2 is several tens of thousands of years .

Table 2. Stages in the weathering of minerals in the < 2 mm fraction of soils (modified table, after White, 1987).

Stage / Type of mineral Soil characteristics
Early weathering stages:

Gypsum (CaSO4* 2H 2O)

Calcite (CaCO3)



Hornblende (amphibole)

Biotite (mica)


These minerals occur in the silt and clay fraction of young soils all over the world, and in soils of arid regions, where lack of water inhibits chemical weathering and leaching. Soils show a very low content of water and organic matter, there is a reducing environment, very limited leaching, and a limited time for weathering.
Intermediate weathering stages:


Hydrous mica (illite)

Vermiculite and mixed layer minerals



Soils found mainly in the temperate regions of the world, frequently on parent materials of glacial or periglacial origin; generally fertile, with grass or forest as the natural vegetation. There is ineffective leaching and cations such as Na, K, Ca, Mg, Fe, and silica are retained.
Advanced weathering stages:


Aluminium oxides (gibbsite)

Iron oxides (goethite, hematite)

Titanium oxides (anatase, rutile, ilmenite)

The clay fractions of many highly weathered soils on old land surfaces of humid and hot intertropical regions are dominated by these minerals. The cations Na, K, Ca, Mg, Fe, and silica are removed from the topsoil due to leaching. Secondary minerals are formed in an oxidizing environment with a low pH where acidic compounds are formed and silica is dispersed.

In the drier and cooler regions of North America 2: 1 type clays tend to dominate soils because of limited weathering. For example, sedimentary or metamorphic rocks containing mica have been an important source of minerals in glacial deposits and therefore those soils are rich in hydrous mica, montmorillonite and vermiculite. If parent rock is sedimentary shale, which is clay rich material, weathering produces 2 : 1, 1 : 1 type clay minerals, or hydrous oxides. Soils with clays high in kaolinite and hydrous oxides tend to be restricted to older landscapes that are both warm and wet. For instance, such soils can be found in the wet and warm climate of the Southern United States. Soils containing hydrous oxides such as Al- and Fe-Oxides as the dominant clays are limited to tropical regions, e.g. soils on the Hawaiian Islands. Parent material high in bases, or a climate which discourage the leaching of bases, encourage montmorillonite formation. For this reason, hydrous mica and montmorillonite are more likely to occur in soils in semiarid and arid climate.

Significance in Pedology

Clay coatings (argillans) are often different in color from and with a higher reflectance than the S-matrix of the ped. They are easily recognizable in sandy and loamy soils, but difficult to distinguish from slickenside surfaces in clay soils. In soil horizons the accumulation of silicate clay is denoted by a ‘t‘ , that are clay coatings on ped faces and / or in pores. The clay coats may be either formed by illuviation or concentrated by migration within the horizon. If slickensides are present, which are formed by shear failure as clay minerals swell upon wetting (vertic charactersitics) the donation is a ‘ss‘.

White R.E., 1987. Introduction to the Principles and Practice of Soil Science. Blackwell Scientific Publ. Inc.

Web Link
The virtual museum of minerals and molecules

Oxides and Hydroxides

General Considerations

Many rocks, primary and secondary minerals contain ions such as silica, iron, aluminum, manganese, and / or smaller amounts of titanium. Oxides and hydroxides may be present as primary minerals (i.e., inherited from the parent material) or secondary / pedogenic minerals (i.e., formed as a result of soil genesis) in soils, whereas several processes are important to consider:

  • Release of metal ions from minerals
  • Translocation of metal ions
  • Transformation processes, such as oxidation-reduction and complexation.

The release of ions such as Si4+, Al 3+, Fe2+, Mn 2+ out of minerals may occur due to protonation or oxidation. In the presence of protons (e.g. protons resulting from a reaction of CO 2 + H2O <- -> H + + HCO3) the silicate may break down by the following reaction (protonation) :

-SiO-Fe2+ + 2H+ <--> Fe2+ + 2HO-SI
(Fe2+ silicate <--> Fe 2+ oxide)

The liberated Fe2+ may then be oxidized immediately after release if the environmental conditions are aerobic (oxygenated conditions), or it may migrate until it reaches an oxygenated zone or remain in a reduced state. The oxidation may occur (i) partly inside the minerals, which results in the release of the oxidized ion out of the mineral because the mineral becomes unstable, or (ii) in solution.

For example, the oxidation of Fe2+ (ferrous ion) to Fe3+ (ferric ion) can be written as:

Fe2+ + 3H2 O = Fe(OH)3 + 3H + + e
(oxidation: release of electrons)

Once released, iron hydrolyzes when it comes in contact with H 2O. The tendency for Fe2+ to hydrolyze and form hydroxides results from two characteristics of iron, i.e., its high affinity for the OH ligand, making the hydrated iron cation a strong acid, and its ready polymerization once hydrolysis proceeds. The resulting iron hydroxides have a low solubility and thus in a pH range > 3 are quite stable in terms of reverse hydrolysis. However, they are vulnerable to transformation in response to increasing reducing conditions:

FeOOH + e + 3H + <--> Fe2+ + 2H 2O
(reduction: acceptance of electrons; remark: generally, reduced species are very mobile)

This occurs (i) whenever oxygen becomes limited (because it has been used by aerobic microorganisms, or induced by water saturation, i.e. anaerobic environmental conditions), (ii) there is a source of organic matter and (iii) appropriate environmental conditions suitable for anaerobic microorganisms to facilitate electron transfer to Fe3+ as part of their metabolic process. Redox reactions are facilitated in soils by the activity of microorganisms (catalysis).

For a given redox couple, the position of the equilibrium depends on the locally prevailing value of pe, which is the negative logarithm of the electron activity, compared to pe0 of the redox couple. The latter value expresses the relative electron activity when reacting species are at unit activity. The aqueous electron is a useful conceptual device for describing the redox status of soils. Soil oxidizability can be expressed by the negative log of the free electron activity:

pe = -log(e).

Large pe values favors electron-poor oxidizing species, whereas small pe values favors electron-rich reducing species.

Reduction processes are often expressed in terms of redox potential (Eh) measured in mV.

Eh = 0.059 pe (at 25° C)

Figure 1. Example of the range in redox potential in soils and the location in the redox range where the various electron acceptors are active (modified after Courtesy of R. W. Miller, 1981).

A high redox potential equals to well-aerated environmental conditions and a low redox potential equals to saturated environmental conditions. Saturated soils become depleted of oxygen, because this is rapidly consumed by aerobic organisms and cannot be replenished by diffusion quickly. Then, anaerobic and facultative organisms continue the decomposition process. In the absence of oxygen, other electron acceptors begin to function, depending on their tendency to accept electrons. When flooding occurs the reduction of the remaining oxygen will take place first, followed by the reduction of nitrate, then manganese, iron, sulphate, and carbon dioxide (Figure 1.). The reduction of oxygen occurs by the O2 consumption of aerobic organisms, NO3 serves as a biochemical electron acceptor involving N-organisms that ultimately excrete reduced N, the reduction of Mn can be initiated in presence of NO 3, whereas the reduction of Fe cannot be initiated in presence of NO 3, and sulfate reducing bacteria are involved to reduce SO 4 2-.

The oxides and hydroxides present in soils reflect the pedoenvironmental conditions of soil formation. The parent material, temperature, moisture, organic material, pH, and Eh control the formation of different types of oxides and hydroxides. Because the oxides and hydroxides, particularly of iron and manganese, show different colors they can be used as an indicator for processes of pedogenesis. It should be stressed that there are continual modifications of pedogenic processes acting in soils. Therefore, the soil properties which can be observed in the field, e.g. soil color expressed by the presence or absence of oxides and hydroxides and the distribution of them, also changes. Thus, it is also important to relate the oxides and hydroxides to contemporary processes of pedogenesis or soil formation associated to some previous periods.

Characteristics of oxides and hydroxides in soils are a relative high cation exchange capacity (CEC), which is due to the dissociation of protons from -OH and -OH2 groups of the hydroxides. This is true also for the oxides, which are often associated by -OH and -OH 2. The CEC for oxides and hydroxides is dependent on pH, where a high pH favors the H+ ions to dissociate from the functional groups and to replace the vacant places with cations. The oxides and hydroxides are efficient sorbents and sinks for:

  • Inorganic ions such as silicate, phosphate, and molybdate
  • Organic anions and molecules such as citrate, fulvic and humic acids
  • Cations such as Al, Cu, Pb, V, Co, Cr, and Ni, some of which are essential for plant growth.

In Figure 2 some soil morphological features associated with different redox statuses and drainage conditions are shown.

Figure 2. Idealized representation of soil morphological features associated with wetness.

Significance in Pedology

In horizons with concretions of hard nodules a ‘c‘ is used for iron, aluminium, manganese, or titanium cemented nodules or concretions. A ‘ g‘ denotes gleying indicated by low chroma color (< 2), either total gleying or the presence of gleying in a mottled pattern. In illuvial horizons the accumulation of organic matter with or without sesquioxides (the oxides and hydroxides of iron and aluminium) are denoted by a 'h ‘. If the sesquioxide component contains enough iron so that color value and chroma exceeds 3, ‘hs‘ is used. Generally, the accumulation of iron and the cementation (i.e., more than 90 % of the horizon is cemented) is denoted by ‘sm‘. Residual accumulation of sesquioxides after intense weathering is denoted by ‘o‘.

Courtesy of R.W. Miller. 1981. The Role of Inorganic Redox Systems in Controlling Reduction in Paddy Soils – from data of W.H. Patrick Jr., Proceedings of the Symposium on Paddy Soil, Institute of Soil Science, Academia Sinica, Science Press, Beijing, and Springer-Verlag, New York.

Iron Oxides and Hydroxides

Primary minerals which contain iron are for example biotite, pyroxene, amphibole, and olivine. Iron oxides and hydroxides are formed by protonation and release of Fe ions out of primary or secondary minerals and / or oxidation. Their occurence provides useful information about soil formation.

Another important attribute of Fe is that its cationic charge is sensitive to changes in the redox status of the soil. This may impart clues via soil color about drainage conditions in the soil. Generally, in soils that do not have impeded drainage the majority of the Fe-oxides occur in the Fe 3+ state, which is typically associated with red, yellow or brown subsoil colors. With a progressive increase in impeded drainage conditions subsoil colors reflect an increasing influence of Fe 2+ on soil color. Fe2+ typically imparts a bluish gray color to poorly drained subsoils, often referred to as a gley color. Along this continuum of drainage conditions variations on patterns may be identified (so-called redoximorphic features (RMF)). In soils that are intermediate between well-drained and waterlogged, periodic reducing conditions may result in particular zones within the soil exhibiting variegated color patterns. These patterns consist of intricate combinations of reddish (high chroma RMFs: containing local concentrations of Fe 3+ oxides), whitish or light gray areas (low chroma RMFs: reflecting zones from which Fe has largely been removed) and sometimes bluish gray areas (gleyed zones: reflecting presence of Fe 2+ -oxides). The distinction between low chroma zones, which are local zones of Fe eluviation, and gleyed zones is not always simple in the field. Both conditions are indicative of reducing conditions, but represent differences in the extent of Fe 2+ translocation. Soil color and in particular patterns of redoximorphic features are used as field indicators of drainage conditions. The criteria are developed to suit local conditions and are ideally based on hydrological and chemical measurements, which are calibrated with observed morphology. This is particularly important because the distribution pattern of Fe-oxide colors may not necessarily reflect contemporary processes operating in the soil. In other words, some of the Fe-oxides still present in the soil may have formed under conditions quite different from those operating at present or they may reflect colors of inherited Fe-oxides or other minerals such as glauconite (greenish color that could be confused as a gley feature) that have not been significantly altered as a result of pedogenesis.

Where iron oxides are absence, soil color usually arise from uncoated mineral grains. They may occur evenly dispersed throughout the soil horizons or as concentrations in particular morphological features such as RMFs, nodules, pipestems. Sesquioxides are a term for the oxides of iron and aluminum and sesquioxic coatings (such as ferrans) can be formed by reduction and solution of Fe under anaerobic conditions and their subsequent oxidation and depositon in aerobic zones. If iron oxides (and manganese oxides) become concentrated in a soil horizon they may form cemented layers, called fragipans(denoted by a ‘x‘), which are hard to very hard and brittle when dry. In contrast, zones of Fe depletion are called neoalbans , which may occur in eluviated soil horizons. The term plinthite is used for B or C horizons (denoted by a ‘v‘), which are humus poor and iron rich. The material usually has reticulate mottling of reds, yellows, and gray colors and hardens irreversibly to ironstone hardpans or aggregates with repeated wetting and drying.

Figure 3. Soft and hard accumulations of iron and manganese in soil peds.

Iron oxides and hydroxides are very stable under aerobic conditions, but they become more soluble under anaerobic conditions (low redox potentials). They are able to form metal-organic complexes, where the metal cations are bonded by functional groups such as -COOH, =CO, -OH, -OCH 3, -NH2, -SH to organic compounds resulting in the formation of a ring structure incorporating the metal ion. These complexes are very stable and called chelates.

In Figure 4 a Eh-pH stability diagram for different iron oxides and hydroxides is shown. The diagram can be used to predict when a species may be oxidized or reduced. Reduction or oxidation can occur outside these boundaries, but only when mediated by an organism and at expense of metabolic energy.

Figure 4. Eh-pH stability diagram for iron oxides and hydroxides (Scheffer et al., 1989).

Several different Fe oxides and hydroxides can be distinguished, which differ in their crystal structure and various other properties (e.g. color, solubility, thermalbehavior). The basic unit of iron oxides and hydroxides ia the Fe(O,OH)6 octahedron. The variation between different iron oxides and hydroxides is mainly due to a variation in the arrangement of these octaheda. Following the 6 most widespread Fe oxides and hydroxides are described briefly:

Goethite (alpha-FeOOH): It is the most frequently occuring Fe-oxide in soil and has a characteristic yellowish brown color. In large concentrations it may also appear as dark brown or black. Goethite is found under a broad range of climatic and hydrological conditions and is the thermodynamically most stable of all Fe-oxides. A variety of conditions appear to favor its formation in preference to hematite, which include cool temperatures, moist soil conditions and presence of comparatively large amounts of organic matter. The absence of hematite and abundance of goethite in many soils of cool or temperate regions supports these general observations. Under suitable conditions, it has been suggested that goethite can form from any Fe source.

Hematite (alpha-Fe2O 3): It has a characteristic bright red color and primarily occurs in the better drained soils of warm temperate to tropical areas. It is also found in restricted areas of cool temperate areas either as a primary mineral or in more highly weathered soils, where it probably formed under climatic conditions warmer and drier than those operating at present. Its formation appears to be favored by warm dry conditions and small amounts of organic matter (OM). There are large areas of intertropics that have yellow topsoils (OM-rich) over red subsoils, which often contain yellow rims around root channels. This supports the concept of the OM-antihematic effect, i.e., hematite formation is hindered in the presence of comparatively large amounts of OM. Hematite probably forms from ferrihydrite (5Fe 2O3*9H 2 O) through aggregation, dehydration, and internal structural rearrangement of the tiny ferrihydrite particles. This has been suggested because ferrihydrite has a hematite-like structure, except that it is highly disordered and is hydrated. As such, ferrihydrite is considered a necessary precursor for hematite formation.

Lepidocrocite (gamma-FeOOH): It has a characteristic bright orange color that is evident in soils where it occurs in large concentrations and is not masked by other pigments. It commonly occurs in association with goethite and usually in soils that have restricted drainage. It forms from the oxidation of Fe2+ compounds, which are common in wet soils. High partial pressure of CO 2 appears to favor goethite formation in preference to lepidocrocite. This has been demonstrated in laboratory experiments and is supported by field observations that show its virtual absence in calcareous soils, and preferential formation of goethite adjacent to roots and lepidocrocite concentrated further away in wet soils. Lepidocrocite formation also appears to be favored by slow oxidation rates and small concentrations of aluminium in the soil solution. Lepidocrocite is rarely found in very acid soils, which typically contain aluminium in the soil solution.

Maghemite (gamma-Fe2O 3): It is common in many soils, notably in the tropics and subtropics and varies in color from red to brown. This oxide occurs in soils, especially those derived from basic igneous rocks. Maghemites are commonly concentrated towards the top of the soil profile. Several pathways have suggested to account for their formation in soils, which include:
(i) Oxidation of magnetite
(ii) Dehydration of lepidocrocite
(iii) Transformation by heating of other Fe-oxides between 300 and 4250C in the presence of organic compounds.

Magnetite: It is listed in most textbooks as a primary soil mineral that does not form through pedogenic processes. It occurs in soils as singular irregular black grains and is present in most soils. It is fairly resistant to weathering but may alter slowly to ferrihydrite, goethite or eventually hematite (via ferrihydrite). It has been documented the occurrence of magnetic bacteria in soils that contain minute crystals of magnetite. This the first account of in situ ‘biogenic’ formation of the mineral in soils.

Ferrihydrite (5Fe2O 3*9H2O): It is a common Fe-oxide found in soils and appears reddish brown in color if not masked by other pigments. It has a poorly ordered structure resembling that of hematite and was previously referred to as amorphous ferric hydroxide. Its formation is favored by rapid oxidation of Fe in the presence of large concentrations of OM and/or silicate. As such it occurs in limited amounts in tropical soils that do not have these conditions. Phosphate anions also appear to favor its formation in preference to other Fe-oxides. The presence of ferrihydrite generally indicates that conditions are not favorable to crystal growth. However, it is generally considered to be a ‘young’ Fe-oxide that will on a pedogenic time scale eventually transform to a more stable and more crystalline Fe-oxide (i.e., crystal-inhibiting substances become degraded, translocated etc., allowing for its transformation).

Some other important iron compounds found in soils are:
Siderite (FeCO3): It is found in waterlogged soils, i.e. under reducing environmental conditions. The color is greenish/blue.Vivianite (Fe3[PO 4]2*H 2O): It is found in waterlogged soils, i.e. under reducing environmental conditions. The color is greenish/blue.

Pyrit (FeS): It is found in waterlogged soils, i.e. under reducing environmental conditions. The color of FeS is black.

Table 1. Summary of iron oxides and hydroxides and their pedoenvironments.

Iron oxide / hydroxide Color Pedoenvironments Soils


yellowish brown, dark brown, to black

7.5 YR – 2.5 Y

Wherever weathering takes place All soils with Fe release

(alpha-Fe2O 3)

bright red

10 R – 5 YR

High soil temperature, better drained soils, rapid biomass turnover, high Fe-release rate from rocks Aerobic soils of the tropics and subtropics


bright orange

5 YR – 7.5 YR

(value >= 6)

Anaerobic > aerobic systems, noncalcareous Aquic subgroups in temperate regions

(gamma-Fe2O 3)

red to brown

2.5 YR – 5 YR

Usually a product of fire Mainly tropical and subtropical soils

(5Fe2O3 *9H2O)

reddish brown

5 YR – 7.5 YR

Rapid oxidation in humic environments wet soils

Scheffer F., and Schachtschabel P. 1989. Lehrbuch der Bodenkunde. Enke Verlag, Stuttgart.

Aluminum Oxides and Hydroxides

The aluminum oxides and hydroxides have a non-distinctive grayish-white color, which is easily masked in soils except when large concentrations occur. In acid soils, deposits of amorphous aluminium hydroxide form in the interlayers of expanding lattice clays, and occur as surface coatings on clay minerals generally. This amorphous material slowly crystallizes to gibbsite (gamma-Al(OH)3), the principal aluminium hydroxide in soil, which is very stable material. This may become dehydrated to form boehmite (alpha-AlOOH), which is common in bauxite deposits but less common in soils. Gibbsite accumulates in old soils that are in an advanced stage of weathering and in younger soils of the tropics.

Al3+ in solution hydrolyzes to produce H+ as follows:

Al3+ + H2 O <--> Al(OH)2+ + H +

The hydroxy aluminium ion may also hydrolyze:

Al(OH)2+ + H2 O <--> Al(OH)2 +1 + H+

Thus, the major source of H+ in moderately and strong acidic soils is aluminium hydrolysis. The Al 3+ cation has a higher charge than other cations such as K +, Na+, Ca 2+, or Mg2+, and in soils with pH <5 the cations associated with minerals are replaced by Al 3+ rather than by H3O +. The ability to replace cations is smaller for H 3O+ in comparison to Al 3+.

Manganese Oxides and Hydroxides

The weathering of primary minerals, such as biotite, pyroxene, amphibole, containing Mn2+, produce in an aerobic environment brown/black Mn4+. The reaction can be written as:

Mn2+ + H2 O = MnO2 + 4H+ + 2e

The pyrolusite (MnO2) is a very stable manganese oxide. Often manganese is associated with other ions such as Ba, Ca, K, Na, Li, NH4, Co, Cu and Ni, therefore the manganese oxides and hydroxides have variable forms. For example, birnessite (Na,Ca,K,Mg,Mn2+ )Mn 64+ O 14 *H2 O, lithiophorite (LiAl 2Mn 2+ Mn 2 4+ O9 *3H 2 O), or hollandite (BaMn 8 O 16). Manganese oxides show even a greater tendency than iron oxides to occur in concretions. A reason might be the reduction of Mn 4+ to Mn 2+, which is relatively soluble, more readily soluble than for example Fe2+ . Manganese-rich micromorphological zones in peds that are black, often also contain large amounts of iron oxides. However, iron-rich concretions have been shown to be low in manganese oxides. Manganese oxide minerals have a black color, which is sometimes difficult to distinct from the black color or organic material.
Figure 5. Eh-pH stability diagram for manganese oxides and hydroxides (Scheffer et al., 1989).

Scheffer F., and Schachtschabel P. 1989. Lehrbuch der Bodenkunde. Enke Verlag, Stuttgart.

Other Non-Silicates

Sulphates: Gypsum (CaSO4 *2H 2O) which is common in soils (-> fertilizers) and jarosite (KFe3 (OH)6 (SO4 )2 which is rarely found in soils are sulphates. An accumulation of gypsum in a B or C horizon is denoted with a ‘y‘. If there is a cementation of more than 90 % and roots can penetrated only through cracks, the symbol ‘ym‘ is used.

Chlorides: The chloride halite (NaCl) is found in soils formed in marine or lacustrine deposits. The accumulation of sodium is denoted by a ‘n‘, indicating a high level of exchangeable sodium in the horizon.

Carbonates: Primary minerals which contain Ca 2+ are calcite (CaCO3 ), dolomite (Ca,Mg)CO3, plagioclase, pyroxene, and amphibole. They are easily weatherable and Ca2+ can be released. The solubility of Ca2+ is dependent on the CO 2 partial pressure (pCO2 ). The higher the pCO2 the higher the pH (pH = -0.67 lg pCO2 + 7.23, where pCO 2 is in kPa) A pH < 7 will be reached if all carbonate is dissolved and begins to leach or migrate. Carbonates have a characteristic to stick mineral particles and organic compounds together to form aggregates. Therefore, carbonates improve soil structure, particularly in A and B horizons. An accumulation of carbonates, usually calcium carbonate, in a horizon is designated with a 'k‘. If more than 90 % of the horizon is cemented by calcium and roots can penetrate only through cracks, the symbol ‘km ‘ is used. If carbonates are leached from the A or even B horizon those layers become acid to very acid. This alters the soil structure, cation exchange capacity, biological activity, etc.

Phosphates: A primary mineral which contains P is apatite (Ca 4(CaF or CaCl)(PO 4) 3. Apatite is stable in non acid soils but when the pH drops below 7 apatite is weathered fast. Orthophosphates (H 2PO 4and HPO 42- ) that are released by mineralization is rapidly adsorbed by the soil particles. This process is called phosphate fixation because the process is difficult to reverse. Phosphate on surfaces that can be readily desorbed and phorphate in solution is called labile phosphate. In contrast, the P held in insoluble compounds or organic matter is called non-labile phosphate (P occluded in surface oxides and insoluble P compounds). Organic P, which is the major source of P for the soil microorganisms and mesofaunas is rapidly mineralized and / or immobilized by the microorganisms, especially bacteria, which have a relative high P requirement. Bacterial phorphate residues comprise mainly the insoluble Ca, Fe, and Al salts of inositol hexaphosphate, the phytates.

Soil Organic Matter (SOM)


Soils represent a major pool (172 x 1010 t) in the cycling of C from the atmosphere to the biosphere and are the habitat for terrestrial photosynthetic organisms, which fix 11 x 10 10 t C per year, about half of which eventually finds its way into soils. Organic matter in soils is represented by plant debris or litter in various stages of decomposition through to humus and includes the living organisms in the soil. Above ground plants (phytomass) are generally excluded from discussions of soil organic matter, but living roots are generally included.

The following definitions will be followed:

  • Soil Organic Matter: Natural C-containing organic materials living or dead, but excluding charcoal.
  • Phytomass: It is the above ground portion of materials of plant origin usually living, but may also include standing dead trees.
  • Litter: It comprises the dead plant and animal debris on the soil surface.
  • Microbial Biomass: It is the living population of soil microrganisms.
  • Macroorganic Matter: Organic fragments from any source which are > 250µm (generally less decomposed than humus).
  • Organic Carbon: The carbon content is commonly used to characterize the amount of organic matter in soils. Organic matter = 1.724 * percent organic carbon.
  • Humus: Material remaining in soils after removal of macroorganic matter (generally material that has been more extensively physically and/or biochemically transformed as a result of soil forming processes than macroorganic matter). There are two major classes: the nonhumic substances (e.g. amino acids, lipids, carbohydrates) and humic substances (a series of high-moelcular-weight amorphous compounds).
  • Humic Acids: Dark-colored amorphous materials that can be extracted from the soil by a variety of reagents, such as strong bases or neutral salts and that are insoluble in dilute acid. This implies that humic acids contain primarily acidic functional groups, such as phenolic or carboxylic groups. Humic acids are composed of molecules with molecular weights in the range 20,000 to 1,360,000. They are considered to be polymerization products of fulvic acids and other decay products.
  • Fulvic Acids: The organic materials that are extracted with humid acid but remain in solution upon acidification with dilute acid. This implies that fulvic acids contain acidic functional groups since it is soluble in strong bases and extracted with humic acids that fulvic acids also contain basic groups since it remains in solution upon acidification. Fulvic acids are composed of molecules with molecular weights in the range 275 to 2,100. They are considered to be decay products of higher plants and microbial residue.
  • Humin: The strong base insoluble fraction.

The carbon cycle describes how carbon is circulated through the atmosphere, biosphere, pedosphere, and hydrosphere. The dead organic matter of the soil is colonized by (micro)organisms, which derive energy for growth from the oxidative decomposition of complex organic molecules. Decomposition is the biochemical breakdown of mineral and organic materials. During decomposition, anorganic elements are converted from organic compounds, a process called mineralization. For example, organic-N and -P is mineralized to NH 4+ and H 2PO4, and C is converted to CO2. The remainder of the substrate C used by the microorganisms is incorporated into their cell substance (biomass), which is called immobilization. The incroporated minerals are immobilized and released after the organisms die or decay. Humification is the formation of humus (complex organic polymers) from raw organic materials, such as fulvic acids, humic acids, or humin.

Figure 1. Transformation of soil organic matter within soil.

Figure 2. The carbon cycle: quantities and reservoirs. (units: 10 9 metric tonnes)

Global Climate Change
Soil organic matter represents a major pool of carbon within the biosphere, estimated at about 1400 x 1015 g globally, roughly twice that in atmospheric CO2. SOM may act as both a source and sink of carbon during global environmental change. Changes in climate are likely to influence the rate of accumulation and decomposition of carbon in SOM, both directly through changes in temperature and soil moisture, and indirectly through changes in plant growth and rhizodeposition. Other factors, especially changes in land use and management, may have even greater effects. Changes in land use or management may occur as a direct result of climate change or other environmental factors, or may be influenced by agricultural, economic or social policies.

SOM models (e.g. DAISY, RothC, CANDY, DNDC, CENTURY, and NCSOIL) embody our best understanding of soil carbon dynamics and may be used to predict how global environmental change will influence soil carbon, and to evaluate the likely effectiveness of different mitigation options.

Factors Influencing SOM

The content of organic matter content is a function of the soil forming factors. Jenny (1930) found that for loamy soils in the United States the effect of soil forming factor to OM were in the order:

  • climate
  • vegetation
  • topography, parent material
  • time

Climate, i.e., precipitation and temperature, influence the amount and type of vegetation as well as the rate of decomposition. The organic matter content of a soil increases with increasing decomposition up to the limit set by temperature. In Figure 3. it is shown, that SOM increases from the east side of the Rocky Mountains to the east with increasing precipitation, and from south to north with decreasing temperature. In soils, every 10 oC increase in mean annual temperature results in the organic matter content being reduced by about 1/3 to 1/2, if all other factors are constant.For example, the carbon content to a depth of 1 m is 2 kg/m 2 in the Badlands of North Dakota, 21 kg/m 2 in some poorly-drained, fine-textured soils of eastern North Dakota. Organic soils (Histosols) contain about 75 kg/m 2(Franzmeier et al., 1985).

Figure 3. Distribution of soil organic matter in the United States, as related to climate and vegetation (adapted from Schreiner et al., 1938).

Generally, cold and arid climate tends to slow down the microbial processes within soil, in particular decomposition and mineralization. Therefore, those soils contain large portions of organic matter as plant debris (macroorganic matter) than as humus. The same effect is observed in acid to very acid soils. The warmer the climate the higher the rates of microbial processes, i.e., the lower the organic matter content in those soils.

The soil moisture content also has a remarkable effect of soil organic matter decomposition and accumulation. Waterlogged soils tend to accumulate organic matter because the microbial processes, in particular decomposition and mineralization, are slowed down. In aquic moisture regimes the drainage and soil aeration is poor (anaerobic conditions). Anaerobic oxidation of organic residues is less efficient than aerobic oxidation. If organic matter is accumulated the soil development is towards organic soils (Histosols). Histosols generally form in wet, poorly aerated sites, such as shallow lakes and ponds, depression areas, swamps, and bogs and are the end product of natural eutrophication.

Vegetation/Soil Organisms
Vegetation affects soil organic matter by the type, amount, and placement of the organic residues. The composition of organic matter in soil can be related to the nature of the soil floral and faunal community. When biomass is added to the soil, three general reactions take place:

  • The bulk of the material undergoes enzymatic oxidation with carbon dioxide, water and heat as major products,
  • N, P and S are released and/or immobilized by a series of reactions unique to each specific element,
  • Compounds resistant to further immediate microbial reaction are formed either from compounds in the initial material or by microbial synthesis.

The rates of decomposition, even for simple substrates such as glucose, vary widely due to differnces in water content, temperature, pH and the availability of nutrients such as P and N to support microbial activity. However, the simpler monomers from carbohydrates, proteins, fats, and many polyphenolic materials are decomposed within weeks in soil environments. Polymers (complex compounds) such as hemi-cellulose or cellulose are decomposed more slowly and their resistance to decomposition increase with complexity. It is essential to emphasize that many of the organic compounds found in soils result from in-situ synthesis mediated by microbial processes. Some natural polymers may persist in soils for years:

  • Cellulose because it is crystalline and often encrusted with lignin and thus not readily accessible to microorganisms,
  • Polyphenols in polymers such as humic materials, and waxes, which are both characteristically recalcitrant (i.e., resistant to rapid microbial breakdown).

Decomposition reactions are catalyzed by enzymes. Generally, when the C : N ratio is > 25, net immobilization occurs, whereas at ratios < 25 net mineralization is likely. Classically, organic matter has been characterized via various extraction/fractionation procedures into non-humic (lipids, carbohydrates and other 'simple' organic compounds) and the more complex humic susbstances (humic acids, fulvic acids and humin). These divisions do not align well with current understanding of the biological and biochemical processes operating during decomposition and stabilization of organic material in soil. Generally, litter from coniferous trees, such as pine, are undergoing a slow decomposition, whereas the litter from deciduous trees, such as elm, ash, oak, and birch, are easy to decompose. Lignin (complex phenolic polymer) is a significant proportion of straw and coniferous litter, which takes a long time to decompose. Coniferous litter tend to be acidic and low in bases, which promotes greater amounts of soil weathering. Annual species, such as grasses, tend to add organic residues not only to the surface, but due to death and decay of the roots. Also, the residue from annual species tends to have higher base contents than are found with perennials. Therefore, a thicker, darker A horizon is formed under grass than under deciduous or coniferous forest. A sequence for decomposition of litter would look like this, whereas it starts with low decomposition and ends with high rates of decomposition:

  • coniferous trees
  • straw
  • deciduous trees
  • grass

Figure 4. Organic matter content in a grassland and a forest soil profile (modified after Foth et al., 1984).

In Figure 4. the differences in organic matter content in a grassland and a forest soil profile are shown. Grassland soils contain more SOM than forest soils under similar environmental conditions. The distribution of SOM is more uniformly distributed through the grassland profile than in a forest soil.

On agricultural land the application of mineral fertilizers, manure or the practice of green manuring influence the organic matter content in soils. The application of manure tends to increase soil organic matter because of the supply of nutrients and organic material to the soil .

Most of the soil organisms are concentrated in the top 15 – 25 cm of soil because C substrates are more plentiful there. Estimates of microbial biomass C range from 500 to 2,000 kg/ha to 15-cm depth. The macro- and mesofaunal biomass ranges from 2 to 5 t/ha, with earthworms making the largest single contribution. Microorganisms use litter and other organic compounds for respiration, where organic material is mineralized and CO2 and inorganic elements are released. The prokaryotes include the bacteria and actinomycetes, the eukaryotes include the fungi, algae and protozoa. They can be classified in heterotrophs, which require C in the form of organic molecules for growth, and the autotrophs, which can synthesize their cell substance from the C of CO2, harnessing the energy of sunlight (in the case of photosynthetic bacteria and algae) or chemical energy from the oxidation of inorganic compounds (the chemoautotrophs). Another way of subdividing the microorganisms is on their requirement of O 2: (i) the aerobes, those requiring O 2 as the terminal acceptor of electrons in respiration (ii) the facultative anaerobes, those normally requiring O 2 but able in anaerobic conditions to use NO 3 and other inorganic compounds as electron acceptor in respiration (iii) the obligate anaerobes, those which grow only in the absence of O 2.

Table 1. Annual rate of litter return to the soil (White, 1987).

Land use / Vegetation type Organic C [t/ha]
Alpine and arctic forest 0.1 – 0.4
Arable land 1.0 – 2.0
Temperate grassland 2.0 – 4.0
Coniferous forest 1.5 – 3.0
Deciduous forest 1.5 – 4.0
Tropical rainforest 5.0 – 10.0

Topography affects the amount of surface runoff, erosion and deposition. If erosion removes soil from the shoulder or backslope areas of a hillslope, thinner and light-colored soils remain where the organic matter content is low. Soils found on footslope or toeslope areas generally show a higher organic matter content and thicker A horizon. Because soil moisture often differs across a hillslope microbial activity is affected as well. For example, north-facing slopes are generally wetter and soil temperature is lower compared to south-facing slopes, therefore the humus content is higher in north and lower in the south facing slopes.

Parent Material
On sandy soils less organic matter is found than on silty or clayey soils. This can be explained by the characteristics of different sized particles. Sandy soils are well aerated and tend to have a low soil moisture content, which are environmental conditions favor for low organic matter content. Vice versa, clayey soils are less aerated with a high amount of fine micropores and tend to have a higher soil moisture content than medium and fine textured soils, hence, they tend to have a high organic matter content. Furthermore, calcareous or Al/Fe rich soils tend to have higher organic matter contents.

“Turnover times” for organic C in soils can be derived by dividing the organic matter content of the soil by the annual biomass input and expressing the answer in years. The turnover time for global C is 30 to 40 years, but varies by orders of magnitude for different ecosystems (the estimates are gross averages and subject to error). Organic soils (Histosols) whose formation is favored by water logging may have turnover times exceeding 2000 y and soils of tundra regions where low temperatures retard oxidation may have turnover times exceeding 100 y. In contrast, the shortest turnover times of about 4 y apply to equatorial forests. Although net primary production is at a maximum in these ecosystems, rapid decomposition precludes appreciable accumulation of soil organic matter.

Foth H.D., 1984. Fundamentals of Soil Science. John Wiley & Sons, New York.

Franzmeier D.P., Lemme G.D., and Miles R.J., 1985. Organic Carbon in Soils of North Central United States. Soil Sci. Soc. Am. J., 49: 702 – 708.

Jenny H., 1930. A Study on the Influence of Climate upon Nitrogen and Organic Matter Content of Soil, Missouri Agr. Exp. Sta. Res. Bull. No 52, University of Missouri, Columbia, Mo.

Schreiner O., and Brown B.E., 1938. Soil Nitrogen. In: Soils and Men. USDA Yearbook, Washington D.C.: 361 – 376.

White R.E., 1987. Introduction to the Principles and Practice of Soil Science. Blackwell Sci. Publ., Oxford, London, Boston.


Cation Exchange Capacity
Organic matter makes a substantial contribution to the cation exchange capacity (CEC) of the whole soil, and hence to the retention of exchangeable cations. This is because humification produces organic colloids of high specific surface area. It should be stressed, that the CEC of soil organic matter is completely pH-dependent and buffered over a wide range of H + ion concentrations. The functional groups, such as the -COOH (carboxylic) and the -OH (phenolic groups), dissociate H + and thus can accept cations such as K +, Na+, Ca 2+ ,or Mg2+. These cations are generally considered to be part of a reservoir of exchangeable cations in the soil. The approximate CEC of organic matter varies between 1500 – 5000 cmol/kg. From 7 – 20 % of the CEC of many soils is caused by organic matter.

Interaction of SOM with Clay-Size Material
The relationship between clay type and content and organic matter accumulation and stabilization is complex. This is because clay content is usually correlated with other factors that result in organic matter production. In particular, clay content is often correlated with greater plant growth for chemical (plant nutrients) and physical (water regime) reasons and results in greater annual input of C. There is also evidence that clay type and associated cations influence organic matter stabilization. Vice versa, the presence of organic matter is of great importance in the formation and stabilization of soil structure. The fulvic and humic acids and their polymers are adsorbed on to mineral surfaces by the functional groups, of which the most important ones are carboxyl (-COOH), carbonyl (-C=O), hydroxyl (-OH), amino (=NH), and amine (-NH 2). Large uncharged polymers (e.g. polysaccharides) can be adsorbed by hydrogen-bonding and by van der Waals’ forces, and also function as bonding agent between mineral particles.

Field and laboratory experiments using additions of 14C-labelled organic compounds have been conducted to evaluate the fate of organic additions to soils of contrasting textures. The finer textured soils typically show a larger initial flush of microbial activity that is followed by greater incorporation and stabilization of organic matter in the soil than found in coarser textured soils.

Porosity exerts a strong influence on the fate of residues added to the soil because it define the domains in which microorganisms can function and those smaller domains into which organic molecules can migrate and become physically isolated from microbial attack. According to Kilbertus (1980), bacteria function only in pores that are at least 3 times their own diameter. Thus, bacteria are excluded from much of the pore space in soils, an effect that becomes more pronounced with increasing clay content. Thus, in clay-rich soils the physical separation between microorganisms and organic molecules can be extensive and account in part for their tendency to have larger accumulations than coarser-textured soils formed under otherwise comparable conditions.

It has been suggested that stabilization of organic molecules may occur between quasi-crystals (a packet of several layers) and within interlayers of 2:1 swelling clays such as montmorillonite. This mechanism has been inferred from examination of high resolution transmission electron micrographs that show presence of organic molecules within ~1.0 µm diameter pores between clay crystals. It is assumed that these domains provide considerable protection against microbial attack. Humic substances coat, partially or totally, mineral particles such as clay, often protecting the coated particles from weathering.

Cation Bridges and Retention of Organic Matter
Polyvalent cations (e.g., Ca2+, Mg 2+, Fe3+, Al 3+) play a major role in the stabilization of organic and inorganic colloids – when in abundance limiting their ability to shrink and swell – favoring a flocculated (stable) condition. Polyvalent cations serve as bridges between negatively charged clays (inorganic colloids) and negatively charged organic colloids, which enhances structural stability.

In neutral and alkaline soils, Ca2+ and Mg2+ are the major cations responsible for bridging and the hydroxypolyvalent cations, Fe3+ and Al3+, serve a similar role in acid soils and those with a large amount of hydrous oxides. There are empirical observations that calcareous soils tend to have larger accumulations of organic matter than their non-calcareous neighbors. Liming experiments provide some insight into the role of Ca2+ in conversion of plant residue into stable organic matter. Addition of CaSO 4 or CaCO3 to soil containing 14C-labelled wheat straw produces an initial ‘priming effect’ on microbial biomass activity resulting in accelerated release of CO 2 that is followed by a greater retention and stabilization of organic matter than found in control treatments (i.e., no Ca 2+ addition). Thus, the ‘priming effect’ of Ca 2+ addition to the soil appears to be transient and the long term effect is one of stabilization of organic matter. The proposed mechanism of stabilization is the formation of Ca2+ cation bridges.

The mechanisms that control Fe3+ and Al3+ linkages with organic molecules are poorly understood. Fe3+ is only sparingly soluble in most soils and occurs mainly in hydrous oxide forms, some of which may be positively charged at low pH because of protonation or addition of hydrogen ions to surface exposed hydroxyl groups. Such positive charged surfaces may attract negatively charged organic molecules. A similar generalized mechanism probably operates with hydrous oxides of Al 3+ . However, at low pH soils may exhibit Al 3+ toxicity to vegetation, which would tend to limit C inputs into the soil.

The chelation process results in the formation of chelates, which are stable complexes containing organic compounds and metallic cations, which are trapped within the ring structures. The complexes can hardly be dissolved. Chelates formed with certain di- and polyvalent cations are the most stable, the stability falling in the order Cu > Fe = Al > Mn = Co > Zn.

Soil Moisture
Humic and fulvic acids are considered to be hydrophilic colloids. As such, they have a high affinity for water and are solvated in aquaeous solution.

Organic compounds (organic colloids < 2 micrometer) have the characteristic to increase field capacity because they tend to hydrolize. Generally, organic matter can hold up to 20 times its weight in water. This is important particularly for sandy soils to improve soil moisture conditions during summer seasons, when precipitation is limited and evapotranspiration rates are high. If organic matter becomes dry it is prone to wind erosion and can be transported over wide distances. Soil Temperature
Because of the dark black color of organic compoundsthe adsorption of solar radiation is high and reflection low, therefore soils high in SOM tend to warm up faster than soils low in SOM.

Organic matter exhibits buffering in slightly acid, neutral, and alkaline range. This buffering helps to maintain an uniform reaction in the soil.


Litter Classification
Litter accumulation and its extent of decomposition on the soil surface (O horizons) differs widely among ecosystems and locally within ecosystems. Climatic factors exert a strong influence on the rate of biomass turnover, and the composition of plant debris and mode(s) of its incorporation into the soil influence activity of the fauna and flora involved in the various transformation processes. The following classification of litter layers or O horizons is based on C:N ratios of the plant debris. In general, debris with a large amount of N is associated with large amount of water soluble organic compounds (e.g., amino acids, sugars) and elements such as S and P that stimulate microbial activity and thus initial degradation of the debris.

  • Mull: low C:N <25, species - alder, false acacia, ash, grasses, legumes (N-fixers), ameliorators.
  • Moder: intermediate C:N 30-45, species – oak, beach.
  • Mor: high C:N >60, species – conifers, ericaceous plants, acidifiers.

A hypothetical soil profile under deciduous species could be described as follows: There is a loose litter layer 2 – 5 cm deep under which the soil is well aggregated, porous, dark-brown in color, and has a granular structure. Below there is a deep A (approximately 30 – 50 cm) of a C : N ratio 10 – 15. The litter accumulation would be classified as mull. In contrast, a hypothetical soil profile under coniferous species could be described as follows: The surface litter is thick (5 – 20 cm) and ramified by plant roots and a fungi mycelium. There is a sharp transition between the organic and underlying mineral soil layers. The litter would be classified as mor.

O horizons are described in the field in terms of their relative degree of decomposition using the following subscript designations:

  • [O] a – Highly decomposed OM, rubbed fibre content < 1/6 of the volume.
  • [O] e – OM of intermediate decomposition, rubbed fibre content 1/6 to 2/5 of volume.
  • [O] i – Slightly decomposed OM, rubbed fibre content > 2/5 of volume.

The symbol ‘h‘ is used for the illuvial accumulation of organic matter but only in combination with B horizons. The ‘h’ indicates an accumulation of illuvial, amorphous, dispersible organic matter with or without sesquioxides. The influence of tillage or other cultivation disturbance that mix the surface layer is denoted by a ‘p‘. The symbol ‘p’ is only used with the master horizon A or O, even if the material mixed by cultivation is from an E, B, or C horizon.

Diagnostic Surface Horizons: Epipedons
For the purposes of soil classification, diagnostic horizons have been developed to provide major distinctions among soils. Diagnostic horizons do not necessarily correspond with those described in the field, but are defined on the basis of specific depth limits and/or presence of specific properties. Diagnostic surface horizons are called epipedons, seven of which are recognized (mollic, ochric, mellanic, plaggen, histic, anthropic and umbric). Epipedons are horizons, which formed at the land surface in which rock/sedimentary structure has been replaced by soil structure, and has either been darkened by soil organic matter (SOM) and/or eluviated. Such a horizon may be covered by thin (< 50 cm) alluvial or eolian material without loosing its identity as an epipedon. Note:

  • If a fresh alluvial or eolian cover is > 50 cm thick, then the underlying horizon is considered to be part of a buried soil, which is indicated by the subscript ‘b‘ following the master horizon designation, e.g., C, Ab, Eb…
  • Any horizon may be at the surface following truncation (erosion) of the soil – in such cases, because the freshly exposed subsurface horizon had not formed at the surface, it would not qualify as an epipedon. Thus, some soils will not have a diagnostic epipedon (e.g. colluvial soils formed in closed depressions, reconstruction sites).

A simple key to understanding the distinctions among the seven diagnostic epipedons centers on (i) distinguishing between mineral and organic surfaces, (ii) the thorough understanding of the definition of the mollic epipedon, and (iii) the awareness of the field settings and conditions where the less common epipedons (e.g., umbric, melanic, anthropic, plaggen) are likely to occur.

Figure 5. illustrates the criterion used to distinguish between mineral and organic soil materials. Note that as clay content increases the amount of SOM required to meet the organic soil material designation increases. The rationale behind this reflects (i) the intrinsic influence that particle-size has on SOM stabilization in soil, and (ii) functional behavior of SOM in relation to particle-size (i.e., SOM has a comparatively stronger influence on soil behavior with decreasing particle-size).

Figure 5. Key to the epipedons in Soil Taxonomy.

The following list describes the epipedons and their major characteristics.

Histic Epipedon: The histic epipedon has an aquic condition for some time in most years or has been artificially drained, and either,

consists of organic soil material,which:

  • – is 20 to 60 cm thick and either contains 75% or more (by volume) sphagnum fibers or has a bulk density, moist, < 0.1 g/cm3 ; or
  • – is 20 to 40 cm thick and meets the organic carbon contents shown in Figure 6.

Is an Ap horizon which, when mixed to a depth of 25 cm, has an organic content (by weight) of:

  • – 16% or more if the mineral fraction contains 60% or more clay: or
  • – 8% or more if the mineral fraction contains no clay; or
  • – 8+ (clay % divided by 7.5) % or more if the mineral fraction contains < 60% clay.

Folistic Epipedon:

  • Consists of organic soil material
  • Epipedon saturated for less than 30 days.

Figure 6. Organic matter (carbon) content required for soil horizons of different clay contents to all qualify as organic horizons.

Mollic Epipedon: The mollic epipedon has the following properties:

  • Soil structure is strong enough so that 1/2 or more of the horizon is not massive when dry. Very coarse prisms, with a diameter of 30 cm or more, are included within the definition of massive if there is no secondary structure within the prisms.
  • Color crushed and smoothed has a Munsell value of 3 or less (moist) and 5 (dry), and a chroma of 3 or less (moist). Additional qualifications on these limits are outlined in Keys to Soil Taxonomy (KST).
  • Base saturation is 50% or more by the NH4 OAc method.
  • Organic carbon is either 0.6% or more through out the thickness of the mollic, or 2.5 % or more in layers that exhibit ‘mollic’ colors.
  • Thickness: After mixing the upper 18 cm of the mineral soil it meets the color and structure requirements outlined above. Additional qualifications on these limits are outlined in Keys to Soil Taxonomy (KST).
  • Phosphorous limits: The epipedon has < 250 ppm of P 2O5 soluble in 1% citric acid. This restriction distinguishes the mollic from cultural epipedons that have unusually large contents of P.
  • Soil moisture regime: If the soil is not irrigated, some part of the epipedon is moist 3 months or more (cumulative) per year in 8 out of 10 y, during times when the soil temperature is 5oC or higher.
  • The n value is less than 0.7. Although many soils that have a mollic epipedon are poorly drained, a mollic does not have the same very high water content as sediments that have been continuously under water since deposition (i.e., they have acquired soil structure, which improves internal drainage).

Umbric Epipedon: The requirements for the umbric epipedon are the same for the mollic, except that base saturation is <50%. Anthropic Epipedon: The requirements for the anthropic epipedon are the same for the mollic, except that P2 O5 soluble in 1% citric acid is > 250 ppm.

Plaggen Epipedon: The plaggen epipedon is a cultural surface horizon produced by long continued manuring. Its color depends on the nature of the manure. Commonly it contains artifacts, such as bits of bricks and pottery through out its depth.

Melanic Epipedon: The melanic epipedon is a thick black horizon which contains high concentrations of organic matter, usually associated with short-range-order minerals or aluminium-humus complexes. The intense black color is attributed to the accumulation of organic matter from which “Type A” humic acids are extracted. This organic matter is thought to result from large amounts of gramineous vegetation, and can be distinguished from organic matter formed under forest vegetation by the melanic index. Additional information about the melanic index is outlined in Keys to Soil Taxonomy (KST).

Ochric Epipedon: The ochric epipedon does not meet the requirements of any of the epipedons listed above, but does show signs of surface soil formation (i.e., soil structure, darkening by organic matter).

The umbric epipedon can not be simply be distinguished from the mollic epipedon in the field. A determination of base saturation is required to distinguish the >50% base saturated mollic from the <50% base saturated umbric. The plaggen epipedon and anthropic epipedon are not commonly found and both owe their origin to local human manipulation of the soil. The histic epipedon has large amounts of organic material overlying mineral subsoils. The histic epipedon is not used in reference to the soils that are classified as Histosols. The mellanic epipedon has restricted occurrence and is associated with soils formed in volcanic materials. Table 2. Summary of epipedon names and important characteristics.

Epipedon Name Derivation Important characteristics
Histic histos, tissue (Greek) Thin, organic horizon saturated 30 consecutive days or more, unless drained. If mixed with mineral material, remains very high in organic matter
Plaggen Plaggen, sod (German) Overly thick mollic (> 50 cm) due to long continued manure application
Anthropic anthropos, man (Greek) Like mollic, but with a high phosphorus content due to long period of cultivation and fertilization
Mollic mollis, soft (Latin) Thick, well-structured, base saturation > 50 %, dark-colored mineral soil horizon
Umbric umbra, shade (Latin) Like mollic, but with base saturation < 50 %
Ochric ochros, pale (Greek) Surface mineral horizon that does not meet criteria for other epipedons

Diagnostic Organic Materials
Fibric Soil Material: In an unrubbed condition, fibers compose over 2/3 of the mass, and the material yields almost clear solutions when extracted with sodium pyrophosphate.

Hemic Soil Material: In an unrubbed condition 1/3 to 2/3 of the total mass is composed of fibers (intermediate in decomposition between fibric and sapric).

Humilluvic Material: Illuvial humus that accumulates after prolonged cultivation of some acid organic soils.

Limnic Soil Material: These are organic or inorganic materials deposited in water by the action of aquatic organisms or derived from underwater and floating organisms. Marl, diatomaceous earth, and sedimentary peat (coprogeneous earth) are considered limnic materials.

Sapric Soil Material: In an unrubbed condition, less than 1/3 of the mass is composed of identifiable fibers and produced sodium pyrophosphate extracts with colors lower in value and higher in chroma than 10 YR 7/3.

Soil Morphology


Soil morphology deals with the form and arrangement of soil features. Micromorphology is using micromorphological techniques (e.g. thin sections) and measurements in the laboratory. Field morphology is the study of soil morphological features in the field by thorough observation, description and interpretation. Observations may be refined with the aid of a hand lens. Simple tests are also used in the field to record salient chemical properties (e.g., pH, presence of carbonates). In addition, field observations and measurements may be refined through a range of laboratory analytical procedures that include more sophisticated evaluation of chemical, biological and physical attributes. However, the quality of field description and sampling ultimately defines the utility of any subsequent laboratory analyses. A keen eye that can discern specific features and their relationship to adjoining features coupled with well-calibrated fingers that can distinguish among relative differences in physical properties of soil material are essential and can only be acquired and maintained through practice. In this course we will focus on field morphology.

Field morphology starts with an in situ examination of a soil profile. Field descriptions are organized by subdividing a vertical exposure of the soil (soil profile) into reasonably distinct layers or horizons that differ appreciably from the horizons immediately above and below in one or more of the soil features listed below. The delineation of horizons is necessarily a somewhat subjective processes because changes in soil attributes are often gradational rather than abrupt. Thus, obvious boundaries between horizons are not always apparent and their assignment may require integrated assessment of changes in several attributes before a sensible and defensible delineation can be made. Knowledge of similar soils and a well-defined rationale for the purpose of the description helps considerably in development of systematic criteria for defining and delineating horizons.

The following information is collected for assembling standard profile descriptions:

  • Depth intervals of horizons or layers (measured from the top of the mineral horizon)
  • Horizon boundary characteristics
  • Color
  • Texture
  • Structure, pores
  • Consistence
  • Roots
  • pH, effervescence
  • Special features such as coatings, nodules, and concretions

Differences between horizons generally reflect the type and intensity of processes that have caused changes in the soil. Ideally, we should always be striving in our descriptions to maintain a link between process and morphology. In many soils, these differences are expressed by horizonation that lies approximately parallel to the land surface, which in turn reflects vertical partitioning in the type and intensity of the various processes that influence soil development. However, there are many exceptions to this preferred horizontal organization.


Master Horizons Master horizons (major horizons) are designated by capital letters, such as O, A, E, B, C, and R.
O horizons: They are dominated by organic material. Some O layers consist of undecomposed or partially decomposed litter, such as leaves, twigs, moss, and lichens, that has been decomposed on the surface; they may be on the top of either mineral or organic soils. Other O layers, are organic materials that were deposited in saturated environments and have undergone decomposition. The mineral fraction of these layers is small and generally less than half the weight of the total mass. In the case of organic soils (peat, muck) they may compose the entire soil profile. Organic rich horizons which are formed by the translocation of organic matter within the mineral material are not designated as O horizons.

A horizons: Mineral horizons that formed at the surface or below an O layer, that exhibit obliteration of all or much of the original rock or depositional structure (in the case of transported materials). A horizons show one or more of the following:

  • An accumulation of humified organic matter intimately mixed with the mineral fraction and not dominated by characteristic properties of the E or B horizons or,
  • Properties resulting from cultivation, pasturing or other similar kinds of disturbance.

E horizons: Mineral horizons in which the main feature is loss of silicate clay, iron, aluminum, or some combination of these, leaving a concentration of sand and silt particles and lighter colors. The horizons exhibit obliteration of all of much of the original rock structure.

B horizons: Horizons in which the dominant feature(s) is one or more of the following:

  • An illuvial concentration of silicate clay, iron, aluminium, carbonates, gypsum, or humus
  • Removal of carbonates
  • A residual concentration of sesquioxides or silicate clays, alone or mixed, that has formed by means other than solution and removal of carbonates or more soluble salts
  • Coatings of sesquioxides adequate to give darker, stronger, or redder colors than overlying and underlying horizons but without apparent illuviation of iron
  • An alteration of material from its original condition that obliterates original rock structure, that form silicate clay, liberates oxides, or both, and that forms a granular, blocky, or prismatic structure
  • Any combination of these.

C horizons: Mineral horizons that are little altered by soil forming processes. They lack properties of O, A, E, or B horizons. The designation C is also used for saprolite, sediments, or bedrock not hard enough to qualify for R. The material designated as C may be like or unlike the material form the A, E, and B horizons are thought to have formed.

R Layers: Consolidated bedrock (hard bedrock), such as granite, basalt, quarzite, sandstone, or limestone. Small cracks, partially or totally filled with soil material and occupied by roots, are frequently present in the R layers.

Transitional Horizons
Transitional horizons are layers of the soil between two master horizons. There are two types of transitional horizons: Horizons dominated by properties of one master horizon that also have subordinate properties of an adjacent master horizon. The designation is by two master horizon capital letters:

  • The first letter indicates the dominant master horizon characteristics
  • The second letter indicated the subordinate characteristics

For example, an AB horizon indicates a transitional horizon between the A and B horizon, but one that is more like the A horizon than the B horizon. An AB or BA designation can be used as a surface horizon if the master A horizon is believed to have been removed by erosion.

Separate components of two master horizons are recognizable in the horizon and at least one of the component materials is surrounded by the others. The designation is by two capital letters with a slash inbetween. The first letter designates the material of greatest volume in the transitional horizon. For example A/B, B/A, E/B or B/E.

Subordinate Distinctions Within Master Horizons
Lower case letters are used to designate specific features within master horizons. They are listed in alphabetical order below:

  • a: Highly decomposed organic material. The ‘a’ is used only with the O master horizon. The rubbed fiber content < 17 % of the volume.
  • b: Buried genetic horizon. It is not used in organic soils or to identify a buried O master horizon.
  • c: Concretions of hard nonconcretionary nodules. This symbol is used only for iron, aluminium, manganese, or titanium cemented nodules or concretions.
  • d: Physical root restriction. It is used to indicate naturally occuring or humanly induced layers such as basal till, plow pans, and other mechanically compacted zones. Roots do not enter except along fracture planes.
  • e: Organic material of intermediate decomposition. This symbol is only used in combination with an O master horizon with rubbed fiber content between 17 – 40 % of the volume.
  • d: Frozen soil. The horizon must contain permanent ice.
  • g: Gleying: This symbol is used in B and C horizons to indicate low chroma color (<= 2), caused by reduction of iron in stagnant saturated conditions. The iron may or may not be present in the ferrous form (Fe 2+). The g is used to indicate either total gleying or the presence of gleying in a mottled pattern. It is not used in E horizons, which are commonly of low chroma, or in C horizons where the low chroma colors are inherited form the parent material and no evidence of saturation is apparent.
  • h: Illuvial accumulation of organic matter: Used only in B horizons. The h indicates an accumulation of illuvial, amorphous, dispersible organic matter with or without sequioxide component. If the sequioxide component contains enough iron so that the color value and chroma exceed 3 additionally a s is used (hs). The organosequioxide complexes may coat sand and silt particles, or occur as discrete pellets, or fill voids and cement the horizon (use of m).
  • i: Slightly decomposed organic material. Used only in combination with an O master horizon to designate that the rubbed fiber content is > 40 % of the volume.
  • k: Accumulation of carbonates, usually calcium carbonate. Used with B and C horizons.
  • m: Cementation or induration: Used with any master horizon, except R, where > 90 % of the horizon is cemented and roots penetrate only through cracks. The cementing material is identified by the appropriate letter:
    • km: carbonate
    • qm: silica
    • sm: iron
    • ym: gypsum
    • kqm: both lime and silica
    • zm: salts more soluble than gypsum
  • n: Accumulation of sodium: This symbol is used on any master horizon showing morphological properties indicative of high levels of exchangeable sodium.
  • o: Residual accumulation of sesquioxides.
  • p: Tillage or other cultivation disturbance (e.g. plowing, hoeing, discing). This symbol is only used in combination with the master horizon A or O.
  • q: Accumulation of silica: This symbol is used with any master horizon, except R, where secondary silica has accumulated.
  • r: Weathered soft bedrock: This symbol is only used in combination with the master C horizon. It designates saprolite or dense till that is hard enough that roots only penetrate along cracks, but which is soft enough that it can be dug with a spade or shovel.
  • s: Illuvial accumulation of sesquioxides and organic matter. This symbol is only used in combination with B horizons. It indicates the presence of illuvial iron oxides. It is often used in conjunction with h when the color is =< 3 (chroma and value).
  • ss: Presence of slickensides. They are formed by shear failure as clay material swell upon wetting. Their presence is an indicator of vertic characteristics.
  • t: Accumulation of silicate clay: The presence of silicate clay forming coats on ped faces, in pores, or on bridges between sand-sized material grains. The clay coats may be either formed by illuviation or concentrated by migration within the horizon. Usually used in combination with B horizons, but it may be used in C or R horizons also.
  • v: Plinthite: This symbol is used in B and C horizons that are humus poor and iron rich. The material usually has reticulate mottling of reds, yellows, and gray colors.
  • w: Development of color and structure. This symbol is used for B horizons that have developed structure or color different, usually redder than that of the A or C horizons, but do not have apparent illuvial accumulations.
  • x: Fragipan character: This symbol is used to designate genetically developed firmness, brittleness, or high bulk density in B or C horizons. No cementing agent is evident.
  • y: Accumulation of gypsum. This symbol is used in B and C horizons to indicated genetically accumulated gypsum.
  • z: Accumulation of salts more soluble than gypsum. This symbol is used in combination with B and C horizons.

Note: Arabic numerals can be added as suffixes to the horizon designations to identify subdivisions within horizons. For example, Bt1 – Bt2 – Bt3 indicated three subsamples of the Bt horizon.

Diagnostic Subsurface Horizons
The accumulation of substances such as silica, iron, aluminium, carbonate, and other salts can result in cemented layers, which change the physical, chemical, and biological behavior of the soil. For example, a cemented layer retards percolation and restrict root activity. Furthermore, the availability of nutrients for plant growth is reduced, i.e., the cation exchange capacity is reduced. There are accumulations in the soil which show the enrichment of one substance and / or the depletion of another substance. This can be expressed by diagnostic subsurface horizons, which are listed in alphabetically order below. It should be stressed that some characteristics can be measured only in the laboratory and not in the field.

Agric horizon: It is formed directly under the plow layer and has silt, clay, and humus accumulated as thick, dark lamallae.

Albic horizon: Typically this is a light-colored E horizon with the color value >= 5 (dry) or >=4 (moist).

Argillic horizon: It is formed by illuviation of clay (generally a B horizon, where the accumulation of clay is denoted by a lower case ‘t’) and illuviation argillans are usually observable unless there is evidence of stress cutans. Requirements to meet an argillic horizon are:

  • 1/10 as thick as all overlying horizons
  • >= 1.2 times more clay than horizon above, or:
  • If eluvial layer < 15 % clay, then >= 3 % more clay, or:
  • If eluvial layer > 40 % clay, then >= 8 % more clay.

Calcic horizon: This layer has a secondary accumulation of carbonates, usually of calcium or magnesium. Requirements:

  • >= 15 cm thick
  • >= 5 % carbonate than an underlying layer

Cambic horizon: This subsurface often shows weak indication of either an argillic or spodic horizon, but not enough to qualify as either. It may be conceptually regarded as a signature of early stages of soil development, i.e soil structure or color development. Requirements:

  • Texture: loamy very fine sand or finer texture
  • Formation of soil structure
  • Development of soil color

Duripan: It is a subsurface horizon cemented by illuvial silica. Air-dry fragments from more than 50 % of the horizon do not slake in water or HCl but do slake in hot concentrated KOH.

Fragipan: These subsoil layers are of high bulk density, brittle when moist, and very hard when dry. They do not soften on wetting, but can be broken in the hands. Air-dry fragements slake when immersed in water. Fragipan genesis as outlined in Soil Taxonomy is largely dependent on physical processes and requires a forest vegetation and minimal physical disturbance. Desiccation and shrinking cause develoment of a network of polygonal cracks in the zone of fragipan formation. Subsequent rewetting washes very fine sand, silt, and clay-sized particles from the overlying horizons into the cracks. Upon wetting, the added materials and plant roots growing into the cracks result in compression or the interprism materials. Close packing and binding of the matrix material with clay is responsible for the hard consistence of the dry prisms. Iron is usually concentrated along the bleached boundaries of the prisms. It has also been postulated that clay and sequioxides cements to be binding agents in fragipans.

Glossic horizon: It occurs usually between an overlying albic horizon and an underlying argillic, kandic, or natric horizon or fragipan. Requirements:

  • >= 5 cm thick
  • Albic material between 15% to 85 %, rest: material like the underlying horizon

Kandic horizon: It is composed of low activity clays, which are accumulated at its upper boundary. Clay skins may or may not be present. It is considered that clay translocation is involved in the process of kandic formation, however, clay skins may be subsequentlz disrupted or destroyed by physical and chemical weathering, or they may have formed in situ. Requirements:

  • Within a distance of < 15 cm at its upper boundary the clay content increases by > 1.2 times
  • Abrupt or clear textural boundary to the upper horizon
  • At pH 7: low-activity clays with CEC of <= 16 cmol/kg and ECEC (effective CEC) of <= 12 cmol/kg

Natric horizon: It is a subsurface horizons with accumulation of clay minerals and sodium. Requirements:

  • Same as argillic horizon
  • Prismatic or columnar structure
  • > 15 % of the CEC is saturated with Na+ , or:
  • More exchangeable Na+ plus Mg 2+ than Ca2+

Oxic horizon: Requirements:

  • >= 30 cm thick
  • Texture: sandy loam or finer
  • At pH 7: CEC of <= 16 cmol/kg and ECEC of <= 12 cmol/kg (i.e., a high content of 1:1 type clay minerals)
  • Clay content is more gradual than required by the kandic horizon
  • < 10 % weatherable minerals in the sand
  • < 5 % weatherable minerals by volume rock structure (i.e., indicative of a very strongly weathered material)

Petrocalcic horizon: It is an indurated calcic horizon. Requirements: At least 1/2 of a dry fragment breaks down when immersed in acid but does not break down when immersed in water

Petrogypsic horizon: This is a strongly cemented gypsic horizon. Dry fragments will not slake in H2O.

Placic horizon: This is a dark reddish brown to black pan of iron and / or manganese. Requirements:

  • 2 – 10 mm thick
  • It has to lie within 50 cm of the soil surface
  • Boundary: wavy
  • Slowly permeable

Salic horizon: This is an subsurface horizon accumulated by secondary soluble salts. Requirements:

  • >= 15 cm thick
  • Enrichment of secondary soluble salts such that electrical conductivity exceeds 30 dS/m more than 90 days each year

Sombric horizon: Formed by illuviation of humus (dark bron to black color) but not of aluminium or sodium. Requirements:

  • At pH 7: base saturation < 50 %
  • Not under an albic horizon
  • Free-draining horizon

Spodic horizon: This horizon has an illuvial accumulation of sequioxides and / or organic matter. There are many specific limitations dealing with aluminium, iron, and organic matter content, and clay ratios, depending on wheather the overlying horizon is virgin or cultivated.

Sulfuric horizon: This is a very acid mineral or organic soil horizon. Requirements:

  • pH < 3.5
  • Mottles are present (yellow color: jarosite)


The boundary between the horizons can be described considering the distinctness and topography. Distinctness refers to the degree of contrast between two adjoining horizons and the thickness of the transition between them. Topography refers to the shape or degree of irregularity of the boundary. In Figure 1 examples for several boundaries are shown.
Figure 1. Boundaries between soil horizons.

Table 1. Classification of horizon boundaries.

Distinctness Abbreviation [cm]
Abrupt a < 2
Clear c 2 – 5
Gradual g 5 – 15
Diffuse d > 15

Topography Abbreviation Description
Smooth s Nearly a plane
Wavy w Waves wider than deep
Irregular i Depth greater than width
Broken b Discontinuous

Soil Color

Color reflects an integration of chemical, biological and physical transformations and translocations that have occurred within a soil . In general, color of surface horizons reflects a strong imprint of biological processes, notably those influenced by the ecological origin of soil organic matter (SOM). Soil organic matter imparts a dark brown to black color to the soil. Generally, the higher the organic matter content of the soil, the darker the soil. A bright-light color can be related to an eluvial horizon, where sequioxides, carbonates and/or clay minerals have been leached out.

Subsoil color reflects more strongly in most soils the imprint of physico-chemical processes. In particular, the redox status of Fe and to a lesser extent Mn, strongly influence the wide variation found in subsoil color. Soil color can provide information about subsoil drainage and the soil moisture conditions of a soils. In well aerated soils, Fe3+ is present which give soil a yellow or reddish color. In more poorly drained soils (anaerobic conditions) iron compounds are reduced and the neutral gray colors of Fe 2+ or bluish-green colors of iron sulfides, iron carbonates, or iron phosphates are visible. A black color in the subsoil can be related to an accumulation of manganese.

In arid and semi-arid environments, the influence of soluble salts (carbonates, sulfates, chlorides etc.) may impart a strong influence on soil color. For example, in arid or sub-humid regions, surface soils may be white due to evaporation of water and soluble salts.

Colors associated with minerals inherited from parent materials may also influence color in horizons that have not been extensively weathered. For example, light gray or nearly white colors is sometimes inherited from parent material, such as marl or quartz. Parent material, such as basalt, can imprint a black color to the subsoil horizons.

Table 2. Soil colors associated with soil attributes.

Soil color Soil attributes Environmental conditions
Brown to black (surface horizon) accumulation of organic matter (OM), humus low temperature, high annual precipitation amounts, soils high in soil moisture, and/or litter from coniferous trees favor an accumulation of OM
Black (subsurface horizon) Accumulation of manganese

Parent material (e.g. basalt)

Bright-light Eluvial horizon (E horizon) In environments where precipitation > evapotranspiration there is leaching of sequioxides, carbonates, and silicate clays. The eluviated horizon consists mainly of silica
Yellow to reddish Fe3+ (oxidized iron) Well-aerated soils
Gray, bluish-green Fe2+ (reduced iron) Poorly drained soils (e.g. subsurface layer with a high bulk density causes waterlogging, or a very fine textured soil where permeability is very low), anaerobic environmental conditions
White to gray Accumulation of salts In arid or subhumid environments where the evapotranspiration > precipitation there is an upward movement of water and soluble salts in the soil
White to gray Parent material: marl, quartz

Soil color is usually registered by comparison of a standard color chart (Munsell Book of Colors). The Munsell notation distinguishs three characteristics of the color: hue, value, and chroma.

  • Hue: It is the dominant spectral color, i.e., whether the hue is pure color such as yellow, red, green, or a mixture of pure colors.
  • Value: It describes the degree of lightness or brightness of the hue reflected in the property of the gray color that is being added to the hue.
  • Chroma: It is the amount of a particular hue added to a gray or the relative purity of the hue.

Figure 2. Munsell soil color chart.

The soil colors are given in the order: hue, value, and chroma. For example, 2.5YR 4/2 describes the hue 2.5YR, dark-grayish brown with a value 4 and a chroma of 2. It should be stressed that soil color is dependent on soil moisture, hence if soil color is recorded also the soil moisture conditions have to be described (e.g. soil color dry, soil color wet). In the upper midwest and other humid areas, colors are conventionally recorded moist. This convention may differ in other climatic regimes.

Many soils have a dominant soil color. Other soils, where soil forming factors vary seasonally (e.g. wet in winter, dry in summer) tend to exhibit a mixture of two or more colors. When several colors are present the term mottling or redoximorphic features (RMF) is used. In such a case, several soil colors have to be recorded, where the dominant color is first, following by a description of the abundance, size, and contrast of the other colors in the mottled pattern. Mottling/RMFs are described by three characteristics: contrast, abundance, and size of area of each color.

Redoximorphic features are a color pattern in a soil due to loss (depletion) or gain (concentration) of pigment compared to the matrix color. It is formed by oxidation / reduction of Fe and/or Mn coupled with their removal and translocation or a soil matrix color controlled by the presence of Fe 2+. RMFs are described separately from other mottles or concentrations! Based on the Field Book for Describing and Sampling Soils (Schoeneberger et al., 1998) RMFs are described in terms of kind, color & contrast, quantity, size, shape, location, composition & hardness, and boundary. RMFs occur in the soil matrix, on or beneath the surface of peds, and as filled pores, linings of pores, or beneath the surface of pores.

Mottles are areas of color that differ from the matrix color. These colors are commonly lithochromic or lithomorphic attributes retained from the geologic source rather than from pedogenesis. Mottles exclude RMFs and ped & void surface features (e.g. clay films). Based on the Field Book for Describing and Sampling Soils (Schoeneberger et al., 1998) mottles are described in terms of quantity, size, color & contrast, moisture state, and shape. Example: Few, medium, distinct, reddish yellow moist (7.5YR 7/8), irregular mottles.

However, a variety of other features in a horizon may have colors different from the matrix, such as infillings of animal burrows (krotovinas), clay coatings (argillans) and precipitates of calcium carbonate. In all instances where specific soil features are described, the shape and spatial relationships of the feature (i.e., where is it located, on a ped face, in the matrix…) to adjacent features should be described in addition to its color, abundance, size and contrast.

Table 3. RMFs/mottles in soils are described in term of abundance, size, and contrast.

Abundance Abbreviation % of the exposed surface
few f < 2
common c 2 – 20
many m 20 – 40
very many v > 40
Size Abbreviation Diameter [mm]
fine 1 < 5 mm
medium 2 5 – 15 mm
coarse 3 > 15 mm

Contrast Abbreviation Visibility
faint f difficult to see, heu and chroma of matrix and mottles closely related
distinst d readily seen, matrix and mottles vary 1 – 2 hues and several units in chroma and value
prominent p conspicious, matrix and mottles vary several units in hue, value, and chroma

Soil Texture Classification

Texture refers to the amount of sand, silt, and clay in a soil sample. The distribution of particle sizes determines the soil texture, which can be assessed in the field or by a particle-size analysis in the laboratory. A field analysis is carried out in the following way: a small soil sample is taken, water is added to the sample, it is kneaded between the fingers and thumb until the aggregates are broken down. The guidelines to determine the particle class are as following:

  • Sand: Sand particles are large enough to grate against each other and they can be detected by sight. Sand shows no stickiness or plasticity when wet.
  • Silt: Grains cannot be detected by feel, but their presence makes the soil feel smooth and soapy and only very slightly sticky.
  • Clay: A characteristic of clay is the stickiness. If the soil sample can be rolled easily and the sample is sticky and plastic when wet (or hard and cloddy when dry) it indicates a high clay content. Note that a high organic matter content tend to smoothen the soil and can influence the feeling for clay.

Table 4. Soil texture classes.

Soil texture Abbreviation
Gravel g
Very coarse sand vcos
Coarse sand cos
Sand s
Fine sand fs
Very fine sand vfs
Loamy coarse sand lcos
Loamy sand ls
Loamy fine sand lfs
Sandy loam sl
Fine sandy loam fsl
Very fine sandy loam vfsl
Gravelly sandy loam gsl
Loam l
Gravelly loam gl
Stony loam stl
Silt si
Silt loam sil
Clay loam cl
Silty clay loam sicl
Sandy clay loam scl
Stony clay loam stcl
Silty clay sic
Clay c

A variety of systems are used to define the size ranges of particles, where the ranges of sand, silt, and clay that define a particle class differs among countries. In the U.S. the soil texture is classified based on the U.S.D.A. system, which is used in this course. The classification of particle sizes are the following (units: mm):

clay: < 0.002

silt: 0.002 – 0.05

fine sand: 0.05 – 0.1

x medium sand: 0.1 – 0.5

coarse sand: 0.5 – 1.0 x x

very coarse sand: 1.0 – 2.0 x

gravel: 2.0 – 762.0

cobbles: > 762.0

Soil texture in the field is determined using a texture triangle (Figure 3). For example, a particle size distribution of 33 % clay, 33 % silt, and 33 % sand would result in the soil texture class ‘clay loam’.
Figure 3. Triangular diagram of soil textural classes (USDA triangle).

Particles greater than 2 mm are removed from a textural soil classification. The presence of larger particles is recognized by the use of modifiers added to the textural class (e.g. gravelly, cobbly, stony) (Table 5 and 6 ).

Table 5. Terms for rock fragments.

Shape and size [mm] Adjective
Spherical and cubelike:

2 – 75

2 – 5

5 – 20

20 – 75

75 – 250

250 – 600

> 600


fine gravelly

medium gravelly

coarse gravelly





2 – 150

150 – 380

380 – 600

> 600





Table 6. Modifier for rock fragments.

Rock fragments by volume [%] Adjectival modifier
< 15 no modifier
15 – 30 gravelly loam
30 – 60 very flaggy loam
> 60 extremely bouldery loam

The distinction between a mineral and an organic horizon is made by the organic carbon content. Layers which contain > 20 % organic carbon and are not water saturated for periods more than a few days are classed as organic soil material. If a layer is saturated for a longer period it is considered to be organic soil material if it has:

  • > = 12 % organic carbon and no clay, or
  • >= 18 % organic carbon and >= 60 % clay, or
  • 12 – 18 % organic carbon and 0 – 60 % clay.

Figure 4. Relationship between soil texture and pore size.

Significance of Soil Texture
The fine and medium-textured soils (e.g. clay loams, silty clay loams, sandy silt loams) are favorable from an agricultural viewpoint because of their high available retention of water and exchangeable nutrients. In fine pores the water is strongly adsorbed in pores but not available for plants, i.e. cohesion and adhesion water occupy the micropore space and they are retained in soil by forces that exceed gravity. In medium-sized pores the available water content is high, whereas in macropores water is more weakly held and percolation is high (gravitational water). In silty soils the distribution of macropores, medium-sized, and fine pores is optimal relating to available water content.

Table 7. Pore size distribution in soils different in texture (Scheffer et al., 1989).

Soils different in texture Pore volume [%] Macropores [%] Medium-sized pores [%] Micropores [%]
Sandy soils 46 (+/- 10) 30 (+/- 10) 7 (+/- 5) 5 (+/- 3)
Silty soils 47 (+/- 9) 15 (+/- 10) 15 (+/- 7) 15 (+/- 5)
Clayey soils 50 (+/- 15) 8 (+/- 5) 10 (+/- 5) 35 (+/- 10)
Organic soils 85 (+/- 10) 25 (+/- 10) 40 (+/- 10) 25 (+/- 10)

In general, coarse-textured soils permit rapid infiltration because of the predominance of large pores, while the infiltration rates of finer-textured soils is smaller because of the predominance of micropores. Other factors, like the compaction of the soil, management practices, vegetation, saturation of the soil have also a significant impact on infiltration and have to be considered.

Soil texture has an impact on soil temperature. Fine-textured soils hold more water than coarse-textured soils, which considering the differences in the specific heat capacity results in a slow response of warming up of fine-textured soils compared to coarse-textured soils.

Another issue to address is the effect that with decreasing particle size the surface area increases. Many important chemical and biological properties of soil particles are functions of particle size and hence surface area. For example, the adsorption of cations (nutrients) or the microbial activity are dependent on surface area.

Scheffer F., and Schachtschabel P., 1989. Lehrbuch der Bodenkunde. Enke-Verlag Stuttgart.

Soil Structure


Structure refers to the arrangement of soil particles. Soil structure is the product of processes that aggregate, cement, compact or unconsolidate soil material. In essence, soil structure is a physical condition that is distinct from that of the initial material from which it formed, and can be related to processes of soil formation. The peds are separated from the adjoining peds by surfaces of weakness. To describe structure in a soil profile it is best to examine the profile standing some meters apart to recognize larger structural units (e.g. prisms). The next step is to study the structure by removing soil material for more detailed inspection. It should be stressed that soil moisture affects the expression of soil structure. The classification of soil structure considers the grade, form, and size of particles.

The grade describes the distinctiveness of the peds (differential between cohesion within peds and adhesion between peds). It relates to the degree of aggregation or the develoment of soil structure. In the field a classification of grade is based on a finger test (durability of peds) or a crushing of a soil sample.

The form is classified on the basis of the shape of peds, such as spheroidal, platy, blocky, or prismatic. A granular or crumb structure is often found in A horizons, a platy structure in E horizons, and a blocky, prismatic or columnar structure in Bt horizons. Massive or single-grain structure occurs in very young soils, which are in an initial stage of soil development. Another example where massive or single-grain structure can be identified is on reconstruction sites. There may two or more structural arrangements occur in a given profile. This may be in the form of progressive change in size/type of structural units with depth (e.g. A horizons that exhibit a progressive increase in size of granular peds that grade into subangular blocks with increasing depth) or occurrence of larger structural entities (e.g. prisms) that are internally composed of smaller structural units (e.g. blocky peds). I such a case all discernible structures should be recorded (i.e. more rather than less detail).

The size of the particles have to be recorded as well, which is dependent on the form of the peds.

Table 8. Classification of soil structure considering grade, size, and form of particles.

Grade Abbreviation Description
Structureless 0 No observable aggregation or no orderly arrangement of natural lines of weakness
Weak 1 Poorly formed indistinct peds
Moderate 2 Well-formed distinct peds, moderately durable and evident, but not distinct in undisturbed soil
Strong 3 Durable peds that are quite evident in undisplaced soil, adhere weakly to one another, withstand displacement, and become separated when soil is disturbed

Form Abbreviation Description
Granular gr Relatively nonporous, spheroidal peds, not fitted to adjoining peds
Crumb cr Relatively porous, spheroidal peds, not fitted to adjoining peds
Platy pl Peds are plate-like. The particles are arranged about a horizontal plane with limited vertical development. Plates often overlap and impair permeability
Blocky bk Block-like peds bounded by other peds whose sharp angular faces form the cast for the ped. The peds often break into smaller blocky peds
Angular blocky abk Block-like peds bounded by other peds whose sharp angular faces form the cast for the ped
Subangular blocky sbk Block-like peds bounded by other peds whose rounded subangular faces form the cast for the ped
Prismatic pr Column-like peds without rounded caps. Other prismatic caps form the cast for the ped. Some prismatic peds break into smaller blocky peds. In these peds the horizontal development is limited when compared with the vertical
Columnar cpr Column-like peds with rounded caps bounded laterally by other peds that form the cast for the peds. In these peds the horizontal development is limited when compared with the vertical
Single grain sg Particles show little or no tendency to adhere to other particles. Often associated with very coarse particles
Massive m A massive structure show little or no tendency to break apart under light pressure into smaller units. Often associated with very fine-textured soils.
Size Abbreviation
Very fine vf
Fine f
Medium m
Coarse c
Very coarse vc

Size Angular and subangular blocky structure

[mm] diameter

Granular and crumb structure

[mm] diameter

Platy structure

[mm] width

Prismatic and columnar structure

[mm] diameter

Very fine < 5 < 1 < 1 (very thin) < 10
Fine 5 – 10 1 – 2 1 – 2 (thin) 10 – 20
Medium 10 – 20 2 – 5 2 – 5 20 – 50
Coarse 20 – 50 5 – 10 5 – 10 (thick) 50 – 100
Very coarse > 50 > 10 > 10 (very thick) > 100

Figure 5. Soil structures (Foth, 1984)

The three characteristics of soil structure are conventionally written in the order grade, size, and shape. For example, weak fine subangular blocky structure. The distribution of different particle sizes in a soil influence the distribution of pores, which can be characterized by their abundance, size, and shape.

Table 9. Abundance, size, and shape of pores.

Abundance Per unit area
Few < 1
Common 1 – 5
Many > 5

Size Diameter (mm)
Very fine < 0.5
Fine 0.5 – 2.0
Medium 2.0 – 5.0
Coarse > 5.0

Vesicular approx. spherical or elliptical
Tubular approx. cylindrical or elongated
Irregularly shaped

Significance of Soil Structure

Soil formation starts with a structureless condition, i.e., the structure is single-grained or massive. Soil development also means development of soil structure, which describes the formation of peds and aggregates. Soil structure forms due to the action of forces that push soil particles together. Subsurface structure tends to be composed of larger structural units than the surface structure. Subsoil structure also tend to have the binding agents on ped surfaces rather than mixed throughout the ped.

Climatically-driven physical processes that result in changes in the amount, distribution and phase (solid, liquid, vapor) of water exert a major influence on formation of soil structure. Phase changes (shrinking-swelling, freezing-thawing) result in volume changes in the soil, which over time produces distinct aggregations of soil materials.

Physico-chemical processes (e.g., freeze-thaw, wet-dry, clay translocation, formation/removal of pedogenic weathering products) influence soil structure formation through out the profile. However, the nature and intensity of these processes varies with depth below the ground surface. The structure and hydrological function of plant communities, texture, mineralogy, surface manipulation and topography all serve to modify local climatic effects through their influence on infiltration, storage and evapotranspiration of water.

Biological processes exert a particularly strong influence on formation of structure in surface horizons. The incorporation of soil organic matter is usually largest in surface horizons. Soil organic matter serves as an agent for building soil aggregates, particularly the polysaccharides appear to be responsible for the formation of peds. Plant roots exert compactive stresses on surrounding soil material, which promotes structure formation. Soil-dwelling animals (e.g., earth worms, gophers) also exert compactive forces, and in some cases (e.g., earth worms) further contribute to structure formation via ingestion/excretion of soil material that includes incorporated organic secretions.

Foth H.D., 1984. Fundamentals in Soil Science. John Wiley & Sons, Inc.


Consistence refers to the cohesion among soil particles and adhesion of soil to other substances or the resistance of the soil to deformation. Whereas soil structure deals with the arrangement and form of peds, consistence deals with the strength and nature of the forces between particles. Consistence is described for three moisture levels: wet, moist, and dry. The stickiness describes the quality of adhesion to other objects and the plasticity the capability of being molded by hands. Wet consistence is when the moisture content is at or slightly more than field capacity. Moist consistence is a soil moisture content between field capacity and the permanent wilting point. When recording consistence it is important to record the moisture status as well. Cementation is also considered when consistence is described in the field. Cementing agents are calcium carbonate, silica, oxides of iron and aluminium.

Table 10. Classification of consistence (Buol et al., 1997).

Moisture status Consistence Abbreviation Description
wet Nonsticky wso Almost no natural adhesion of soil material to fingers
Slightly sticky wss Soil material adheres to only one finger
Sticky ws Soil material adheres to both fingers
Very sticky wvs Soil material strongly adheres to both fingers
Nonplastic wpo No wire is formable by rolling material between the hands
Slightly plastic wps Only short (< 1cm) wires are formed by rolling material between the hands
Plastic wp Long wires (>1cm) can be formed and moderate pressure is needed to deform a block of the molded material
Very plastic wvp Much pressure is needed to deform a block of the molded material
Moist Loose ml Soil material is noncoherent
Very friable mvfr Aggregates crush easily between thumb and finger
Friable mfr Gentle pressure is required to crush aggregates
Firm mfi Moderate pressure is required to crush aggregates
Very firm mvfi Strong pressure is required to crush aggregates
Extremely firm mefi Aggregates cannot be broken by pressure
Dry Loose dl
Soft ds
Slightly hard dsh
Hard dh
Very hard dvh
Extremely hard deh
Cementation Weakly cemented cw
Strongly cemented cs
Indurated ci

Buol S.W., Hole F.D., McCracken R.J., and Southard R.J., 1997. Soil Genesis and Classification. Iowa State University Press.


Plant roots give evidence of the plant root activity and the penetration. For example, it is important to record if roots only penetrate through cracks, are retarded by waterlogged layers or cemented layers. Other reasons for limited root penetration can be soil compaction or the absence of nutrients. If there is no obstacle to root growth in the soil the roots may be distributed evenly in a soil. It is important to record the quantity and diameter of roots.

Table 11. Classification of roots.

Root quantity classes Per unit area
Very few < 0.2
Moderately few 0.2 to 1
Few < 1
Common 1 to < 5
Many >= 5

Size classes of roots Diameter in mm
Very fine < 1
Fine 1 – 2
Medium 2 – 5
Coarse 5 – 10
Very coarse > 10

PH and Effervescence

The acidity of a soil can be tested using a simple field test set for fast pH determination. The pH is important for the pH dependent charge of silicates and organic material, therefore for the cation exchange capacity. Furthermore, the pH determines which buffering system is active, i.e. how soils can cope with additional H+ ions. For example, buffering systems are carbonates, organic matter, silicates, or iron and aluminium oxihydroxides.

Using HCl on a small soil sample the reaction (effervescence) can give clues of the calcium carbon content in the sample.

2 HCl + CaCO3 <–> CaCl 2 + H2CO 3 (effervescence)

Special Features

Special features occur is soils which should be recorded additionally. Ped exteriors include clay coats, organic matter coats, silt coats, sand coats, carbonate coats, manganese coats, slickensides, stress surfaces, and clay bridges between sand grains. Ped interiors include concentrations of oxides, nodules, soft accumulations, pseudo-rock fragments, plinthite, and streaks. In particular, concretions are resulting from alternate periods of reducing and oxidizing regimes. Another special feature might be the evidence of animal activity by burrowing animals or high earthworm activity.


Def: Soil features that form by accumulation of material during pedogenesis.

Processes involved: Chemical dissolution/precipitation, oxidation and reduction, physical and/or biological removal, transport, and accumulation


  • Finely disseminated materials: Small precipitates (e.g. salts, carbonates) dispersed throughout the matrix of a horizon
  • Concentrations
  • Masses: Non-cemented bodies of accumulation of various shapes that cannot be removed as discrete units (e.g. crystalline salts)
  • Nodules: Cemented bodies of various shapes that can be removed as discrete units from soil
  • Concretions: Cemented bodies similar to nodules, except for the presence of visible, concentric layers of material around a point, line, or plane
  • Crystals: Macro-crystalls forms of relatively soluble salts (e.g. gypsum, carbonates) that form in situ by precipitation from soil solution
  • Biological concentrations: Discrete bodies accumulated by a biological process (e.g., fecal pellets, insect casts)

Ped & Void Surface Features

These features are coats/films or stress features formed by translocation and deposition, or shrink-swell processes on or along surfaces. They are described in terms of kind, amount, continuity, distinctness, location, and color. Examples: Ferriargillans (Fe 3+ stained clay films) Mangans (black, thin films of Mn)