Cobbles on a beach

River Rhône flowing into Lake Geneva

Sediment billowing out from Italy's shore into the Adriatic

Sediment, a naturally occurring material that is broken down by processes of weathering and erosion, and is subsequently transported by the action of fluids such as wind, water, or ice, and/or by the force of gravity acting on the particle itself.

Sediments are most often transported by water (fluvial processes), wind (aeolian processes) and glaciers. Beach sands and river channel deposits are examples of fluvial transport and deposition, though sediment also often settles out of slow-moving or standing water in lakes and ocean dunes and loess are examples of aeolian transport and deposition. Glacial moraine deposits and till are ice-transported sediments.

Classification[link]

Sediment can be classified based on its grain size and/or its composition.

Grain size[link]

Sediment in the Gulf of Mexico

Sediment off the Yucatán Peninsula

Sediment size is measured on a log base 2 scale, called the "Phi" scale, which classifies particles by size from "colloid" to "boulder".

Composition[link]

Composition of sediment can be measured in terms of:

• parent rock lithology
• mineral composition
• chemical make-up.

This leads to an ambiguity in which clay can be used as both a size-range and a composition (see clay minerals).

Sediment transport[link]

Sediment builds up on human-made breakwaters because they reduce the speed of water flow, so the stream cannot carry as much sediment load.

Glacial transport of boulders. These boulders will be deposited as the glacier retreats.

Sediment is transported based on the strength of the flow that carries it and its own size, volume, density, and shape. Stronger flows will increase the lift and drag on the particle, causing it to rise, while larger or denser particles will be more likely to fall through the flow.

Fluvial processes: rivers, streams, and overland flow[link]

Particle motion[link]

Rivers and streams carry sediment in their flows. This sediment can be in a variety of locations within the flow, depending on the balance between the upwards velocity on the particle (drag and lift forces), and the settling velocity of the particle. These relationships are given in the following table for the Rouse number, which is a ratio of sediment fall velocity to upwards velocity.

Failed to parse (Missing texvc executable; please see math/README to configure.): \textbf{Rouse}=\frac{\text{Settling velocity}}{\text{Upwards velocity from lift and drag}}=\frac{w_s}{\kappa u_*}

where

• Failed to parse (Missing texvc executable; please see math/README to configure.): w_s

is the fall velocity

• Failed to parse (Missing texvc executable; please see math/README to configure.): \kappa

is the von Kármán constant

• Failed to parse (Missing texvc executable; please see math/README to configure.): u_*

is the shear velocity

If the upwards velocity approximately equal to the settling velocity, sediment will be transported downstream entirely as suspended load. If the upwards velocity is much less than the settling velocity, but still high enough for the sediment to move (see Initiation of motion), it will move along the bed as bed load by rolling, sliding, and saltating (jumping up into the flow, being transported a short distance then settling again). If the upwards velocity is higher than the settling velocity, the sediment will be transported high in the flow as wash load.

As there are generally a range of different particle sizes in the flow, it is common for material of different sizes to move through all areas of the flow for given stream conditions.

Fluvial bedforms[link]

Modern asymmetric ripples developed in sand on the floor of the Hunter River, New South Wales, Australia. Flow direction is from right to left.

Sinuous-crested dunes exposed at low tide in the Cornwallis River near Wolfville, Nova Scotia

Ancient channel deposit in the Stellarton Formation (Pennsylvanian), Coalburn Pit, near Thorburn, Nova Scotia.

Sediment motion can create self-organized structures such as ripples, dunes, antidunes on the river or stream bed. These bedforms are often preserved in sedimentary rocks and can be used to estimate the direction and magnitude of the flow that deposited the sediment.

Surface runoff[link]

Overland flow can erode soil particles and transport them downslope. The erosion associated with overland flow may occur through different methods depending on meteorological and flow conditions.

• If the initial impact of rain droplets dislodges soil, the phenomenon is called rainsplash erosion.
• If overland flow is directly responsible for sediment entrainment but does not form gullies, it is called "sheet erosion".
• If the flow and the substrate permit channelization, gullies may form; this is termed "gully erosion".

Key fluvial depositional environments[link]

The major fluvial (river and stream) environments for deposition of sediments include:

• Deltas (arguably an intermediate environment between fluvial and marine)
• Point bars
• Alluvial fans
• Braided rivers
• Oxbow lakes
• Levees
• Waterfalls

Aeolian processes: wind[link]

Wind results in the transportation of fine sediment and the formation of sand dune fields and soils from airborne dust.

Glacial processes[link]

Glacial sediments from Montana

Glaciers carry a wide range of sediment sizes, and deposit it in moraines.

Mass balance[link]

The overall balance between sediment in transport and sediment being deposited on the bed is given by the Exner equation. This expression states that the rate of increase in bed elevation due to deposition is proportional to the amount of sediment that falls out of the flow. This equation is important in that changes in the power of the flow changes the ability of the flow to carry sediment, and this is reflected in patterns of erosion and deposition observed throughout a stream. This can be localized, and simply due to small obstacles: examples are scour holes behind boulders, where flow accelerates, and deposition on the inside of meander bends. Erosion and deposition can also be regional: erosion can occur due to dam removal and base level fall. Deposition can occur due to dam emplacement that causes the river to pool, and deposit its entire load or due to base level rise.

Shores and shallow seas[link]

Seas, oceans and lakes accumulate sediment over time. The sediment could consist of terrigenous material, which originates on land, but may be deposited in either terrestrial, marine, or lacustrine (lake) environments; or of sediments (often biological) originating in the body of water. Terrigenous material is often supplied by nearby rivers and streams or reworked marine sediment (e.g. sand). In the mid-ocean, living organisms are primarily responsible for the sediment accumulation, their shells sinking to the ocean floor upon death.

Deposited sediments are the source of sedimentary rocks, which can contain fossils of the inhabitants of the body of water that were, upon death, covered by accumulating sediment. Lake bed sediments that have not solidified into rock can be used to determine past climatic conditions.

Key marine depositional environments[link]

Holocene eolianite and a carbonate beach on Long Island, Bahamas.

The major areas for deposition of sediments in the marine environment include:

• Littoral sands (e.g. beach sands, runoff river sands, coastal bars and spits, largely clastic with little faunal content)
• The continental shelf (silty clays, increasing marine faunal content).
• The shelf margin (low terrigenous supply, mostly calcareous faunal skeletons)
• The shelf slope (much more fine-grained silts and clays)
• Beds of estuaries with the resultant deposits called "bay mud".

One other depositional environment which is a mixture of fluvial and marine is the turbidite system, which is a major source of sediment to the deep sedimentary and abyssal basins as well as the deep oceanic trenches.

Any depression in a marine environment where sediments accumulate over time is known as a sediment trap.

The null point theory explains how sediment deposition (geology) undergoes a hydrodynamic sorting process within the marine environment leading to a seaward fining of sediment grain size.

Environmental issues[link]

Erosion and agricultural sediment delivery to rivers[link]

One cause of high sediment loads from slash and burn and shifting cultivation of tropical forests. When the ground surface is stripped of vegetation and then seared of all living organisms, the upper soils are vulnerable to both wind and water erosion. In a number of regions of the earth, entire sectors of a country have become erodible. For example, on the Madagascar high central plateau, which constitutes approximately ten percent of that country's land area, most of the land area is devegetated, and gullies have eroded into the underlying soil in furrows typically in excess of 50 meters deep and one kilometer wide.^{[citation needed]} This siltation results in discoloration of rivers to a dark red brown color and leads to fish kills.

Erosion is also an issue in areas of modern farming, where the removal of native vegetation for the cultivation and harvesting of a single type of crop has left the soil unsupported. Many of these regions are near rivers and drainages. Loss of soil due to erosion removes useful farmland, adds to sediment loads, and can help transport anthropogenic fertilizers into the river system, which leads to eutrophication.

Dregs[link]

Sediment in wine, beer, Turkish coffee or other beverages is known as dregs.

See also[link]

• Bar (river morphology)
• Beach cusps
• Biorhexistasy
• Bioswale
• Decantation
• Erosion
• Exner equation
• Particle size (grain size)
• Regolith
• Sand
• Sediment precipitation
• Sediment trap
• Sedimentary depositional environment
• Settling
• Surface runoff
• Deposition (geology)

References[link]

• Prothero, Donald R.; Schwab, Fred (1996), Sedimentary Geology: An Introduction to Sedimentary Rocks and Stratigraphy, W. H. Freeman, ISBN 0-7167-2726-9 
• Siever, Raymond (1988), Sand, New York: Scientific American Library, ISBN 0-7167-5021-X 
• Nichols, Gary (1999), Sedimentology & Stratigraphy, Malden, MA: Wiley-Blackwell, ISBN 0-632-03578-1 
• Cambridge, MA, H. G. (1978), Sedimentary Environments: Processes, Facies and Stratigraphy, Blackwell Science, ISBN 0-632-03627-3 

A represents soil; B represents laterite, a regolith; C represents saprolite, a less-weathered regolith; the bottommost layer represents bedrock

Loess field in Germany

Surface-water-gley developed in glacial till, Northern Ireland

Soil is a natural body that consists of layers (soil horizons), composed primarily of minerals, which differ from their parent materials in their texture, structure, consistency, color, chemical, biological and other physical characteristics.^[1] The end result, soil, is the end product of the influence of the climate (temperature, precipitation), relief (slope), organisms (flora and fauna), parent materials (original minerals), temperature, and time. In engineering, soil is referred to as regolith, or loose rock material. Strictly speaking, soil is the depth of regolith that influences and has been influenced by plant roots and may range in depth from centimeters to many meters.

Soil is composed of particles of broken rock (materials) that have been altered by chemical and mechanical processes that include weathering, erosion and precipitation. Soil is altered from its parent material by the interactions between the lithosphere, hydrosphere, atmosphere, and biosphere.^[2] It is a mixture of mineral and organic materials that are in solid, gaseous and aqueous states.^[3][4] Soil is commonly referred to as earth or dirt; technically, the term dirt should be restricted to displaced soil.^[5]

Soil forms a structure filled with pore spaces that can be thought of as a mixture of solids, water and air (gas).^[6] Accordingly, soils are often treated as a three state system.^[7] Most soils have a density between 1 and 2 g/cm³.^[8] Little of the soil of planet Earth is older than the Tertiary and none older than the Pleistocene.^[9]

Darkened topsoil and reddish subsoil layers are typical in some regions.

On a volume basis a good quality soil is one that is 45% minerals (sand, silt, clay), 25% water, 25% air, and 5% organic material, both live and dead. The mineral and organic components are considered a constant while the percentages of water and air are the only variable parameters where the increase in one is balanced by the reduction in the other.

Given time, the simple mixture of sand, silt, and clay will evolve into a soil profile that consists of two or more layers called horizons that differ in one or more properties such as texture, structure, color, porosity, consistency, and reaction. The horizons differ greatly in thickness and generally lack sharp boundaries. Mature soil profiles in temperate regions may include three master horizons A, B and C. The A and B horizons are called the solum or “true soil” as most of the chemical and biological activity that has formed soil takes place in those two profiles.^[10]

The pore space of soil is shared by gasses as well as water. The aeration of the soil influences the health of the soil's flora and fauna and the emission of greenhouse gasses.

Of all the factors that influence the evolution of soil, water is the most powerful due to its effect on the solution and precipitation of minerals, plant growth, the leaching of minerals from the soil profile and the transportation and deposition of the very materials of which a soil is composed.

Soil colloidal particles (clay and humus) behave as a repository of nutrients and moisture, and buffer the variations of soil solution ions. Their contributions to soil nutrition are out of proportion to their part of the soil. Colloids act to store nutrients that might be leached and to release those ions in response to soil pH.

Soil pH, a measure of the hydrogen ion (acid-forming) soil reactivity, is a function of the soil materials, precipitation level and plant root behavior. Soil pH affects the availability of nutrients.

Most nutrients, with the exception of the lack of nitrogen in desert soils, are present in the soil but may not be available due to high or low pH. Most nutrients originate from minerals, are stored in organic material both live and dead and on colloidal particles as ions. The action of microbes on organic matter and minerals may free nutrients for use, sequester them, or cause their loss from the soil by their volitalization to gasses. The majority of the nitrogen available in soils are the result of nitrogen fixation by bacteria.

The organic material of the soil has a powerful effect on its development, fertility and available moisture. Following water, organic material is next in importance to the formation and fertility of soil.

History of the study of soil[link]

The history of the study of soil is intimately tied to our urgent need to provide food for ourselves and forage for our animals. Throughout history, civilizations have prospered or declined as a function of availability and productivity of their soils.

The Greek historian Xenophon (450-355 B.C.) is credited with being the first to expound upon the merits of green-manuring crops, "But then whatever weeds are upon the ground, being turned into earth, enrich the soil as much a dung."^[11]

Columella’s Husbandry, circa 60 A.D. was used by 15 generations (450 years) of those encompassed by the Roman Empire until its collapse. From the fall of Rome to the French Revolution, knowledge of soil and agriculture was passed on from parent to child and as a result, crop yields were low. During the Dark Ages for Europe, Yahya Ibn_al-'Awwam’s handbook guided the people of North Africa, Spain and the Middle East with its emphasis on irrigation, a translation of which was finally carried to the southwest of the United States.

Jethro Tull, an English gentleman, introduced in 1701 an improved grain drill that systemized the planting of seed and invented a horse-drawn weed hoe, the two of which allowed fields once choked with weeds to be brought back to production and seed to be used more economically. Tull, however, also introduced the mistaken idea that manure introduced weed seeds, and that fields should be plowed in order to pulverize the soil and so release the locked up nutrients. His ideas were taken up and carried to their extremes in the 20th century, when farmers repeatedly plowed fields far beyond what was necessary to control weeds. During a period of drought, the repeated plowing resulted in the Dust Bowl in the prairie region of the Central United States and Canada.

The "two-course system" of a year of wheat followed by a year of fallow was replaced in the 18th century by the Norfolk four-course system, in which wheat was grown in the first year, turnips the second, followed by barley, clover and ryegrass together, in the third. The taller barley was harvested in the third year while the clover and ryegrass were grazed or cut for feed in the fourth. The turnips fed cattle and sheep in the winter. The fodder crops produced large supplies of animal manure, which returned nutrients to the soil.^[12] By the mid-nineteenth century, the Norfolk four-course system was widely adopted throughout Europe.

Experiments into what made plants grow, first led to the idea that the ash left behind when plant matter was burnt was the essential element but overlooked the role of nitrogen, which is not left on the ground after combustion. Jan Baptist van Helmont thought he had proved water to be the essential element from his famous experiment with a willow tree grown in carefully controlled conditions in which only water was added, and after five years of growth was removed and weighed, roots and all, found to be 165 pounds. The oven-dried soil, originally 200 pounds, was again dried and weighed and found to have lost only two ounces, which van Helmont reasonably explained as experimental error and assumed that the soil had in fact lost nothing. As rain water was the only thing added by the experimenter, he concluded that water was the essential element in plant life. In fact the two ounces lost from the soil were the minerals taken up by the willow tree during its growth.

John Woodward experimented with various types of water ranging from clean to muddy and found muddy water the best, and so he concluded earthy matter was the essential element. Others concluded it was humus in the soil that passed some essence to the growing plant.

The French chemist Antoine Lavoisier showed that plants and animals must “combust” oxygen internally to live and was able to deduce that most of the 165-pound weight of van Helmont’s willow tree derived from air. Hence, the chemical basis of nutrients delivered to the soil in manure was emphasized and in the mid-19th century chemical fertilizers were used, but the dynamic interaction of soil and its life forms awaited discovery.

It was known that nitrogen was essential for growth and in 1880 the presence of Rhizobium bacteria in the roots of legumes explained the increase of nitrogen in soils so cultivated. The importance of life forms in soil was finally recognized.

Crop rotation, mechanization, chemical and natural fertilizers led to a doubling of wheat yields in western Europe between 1800 to 1900.^[13]

Soil forming factors[link]

Soil formation, or pedogenesis, is the combined effect of physical, chemical, biological, and anthropogenic processes on soil parent material. Soil genesis involves processes that develop layers or horizons in the soil profile. These processes involve additions, losses, transformations and translocations of material that compose the soil. Minerals derived from weathered rocks undergo changes that cause the formation of secondary minerals and other compounds that are variably soluble in water. These constituents are moved (translocated) from one area of the soil to other areas by water and animal activity. The alteration and movement of materials within soil causes the formation of distinctive soil horizons.

How the soil "life" cycle proceeds is influenced by at least five classic soil forming factors that are dynamically intertwined in shaping the way soil is developed: parent material, climate, topography (relief), organisms and the passage of time. When reordered to climate, relief, organisms, parent material and time they form the acronym CROPT.^[14]

An example of soil development would begin with the weathering of lava flow bedrock which would produce the purely mineral-based parent material from which soils form. Soil development would proceed most rapidly from bare rock of recent flows in a warm climate, under heavy and frequent rainfall. In such a condition, plants become established very quickly on basaltic lava, even though there is very little organic material. The plants are supported by the porous rock as it is filled with nutrient-bearing water that carries dissolved minerals from rocks and guano. Crevasses and pockets, local topography of the rocks, would hold fine materials and harbor plant roots. The developing plant roots themselves are associated with mycorrhizal fungi^[15] that gradually break up the porous lava, and by these means organic matter and a finer mineral soil accumulate with time.

Parent material[link]

The material from which soil forms is called parent material. Rock, whether its origin is igneous, sedimentary or metamorphic, is the source of all soil mineral materials. The formation of a soil is dependent on their transportation and deposition and the physical and the chemical weathering as original minerals are transformed into soil.

Typical soil mineral materials are:^[16]

• Quartz: SiO₂
• Calcite: CaCo₃
• Feldspar: KAlSi₃O₈
• Mica (biotite): K(Mg,Fe)₃AlSi₃O₁₀(OH)₂

Classification of parent material[link]

Parent materials may be classified according to how they came to be deposited in place. Residual materials are those that have been weathered in place from primary bedrock; transported material have been deposited by water, wind, ice or gravity; and cumulose material is organic matter developed and accumulated in place.

Residual soils are soils that develop from their underlying parent rocks and have the same general chemistry as their parent rocks. The soils found on mesas, plateaus and plains are residual soils but few other soils are residual. In the United States as little as three percent of the soils are residual soils.^[17]

Most soils are not residual but derive from transported materials that have been moved many miles by wind, water, ice and gravity.^[18]

• Aeolian processes are capable of moving silt and fine sand many hundreds of miles forming loess soils (60-90 percent silt),^[19] common in the Midwest of North America and in Central Asia. Clay is seldom moved by wind as it forms stable aggregates.

• Water transported material are classed as either alluvial, lacustrine, or marine. Alluvial materials are those moved and deposited by flowing water. Sedimentary deposits accumulated in lakes and are called lacustrine. Lake Bonneville, and many soils around the Great Lakes of the United States are examples. Marine deposits along the Atlantic and Gulf Coasts and in the Imperial Valley of California are deposits from ancient seas that have been revealed as the land uplifted.

• Ice moves parent material and makes deposits in the form of terminal and lateral moraines in the case of stationary glaciers. Retreating glaciers leave smoother ground moraines and in all cases, outwash plains are left as alluvial deposits are moved downstream from the glacier.

• Parent material moved by gravity is obvious at the base of steep slopes as talus cones and is called colluvial material.

Cumulose parent material originates from deposited organic material and includes peat and mucksoils and result from plant residues that have been preserved by the low oxygen content of a high water table.

Weathering of parent material[link]

The weathering of parent material takes the form of physical disintegrating and chemical decomposition and transformation.

• Physical disintegration (weathering), the first stage in the transforming of parent material into soil material, may result from the freezing of absorbed water, causing the physical splitting of material, along a path toward the center of the rock, while temperature gradients can cause exfoliation of "shells" of rock. Cycles of wetting and drying, cause soil particles to grind into finer size as does the physical rubbing of material caused by wind, water and gravity. Organisms, as a result of their roots or digging, also reduce parent material in size.

• Chemical decomposition (weathering) results when minerals are made soluble or are changed in structure. The solution of salts in water results from the action of bipolar water on the ionic salt compound. In other cases the minerals are transformed into polar molecules and then pulled into solution with water. The hydrolysis of orthoclase-feldspar transforms it to acid silicate clay and potassium hydroxide which are more soluble in water. Carbon dioxide in solution with water forms carbonic acid, transforms calcite into calcium bicarbonate which is much more soluble in water. Structural changes to parent material result from hydration, oxidation and reduction. Hydration causes minerals to swell, often leaving them stressed and more easily decomposed. Oxidation of minerals changes their volume and changes the oxidation number of some element in the mineral structure leaving them with a net electrical charge and more prone to attack by water and carbonic acid. The reduction of minerals, which occurs most often in oxygen poor conditions, increases the negative valence of the mineral making it more easily decomposed.^[20]

Saprolite is a particular example of a residual soil formed from the transformation of granite, metamorphic and other types of bedrock into clay minerals. Often called "weathered granite", saprolite is the result of weathering processes that include: hydrolysis (the division of a mineral into acid and base pairs by the splitting of intervening water molecules), chelation from organic compounds, hydration (the solution of minerals in water with resulting cation, anion pairs), and physical processes that include freezing and thawing.^[21] The mineralogical and chemical composition of the primary bedrock material, its physical features, including grain size and degree of consolidation, plus the rate and type of weathering, transforms the parent material into a different mineral. The texture, pH and mineral constituents of saprolite are inherited from its parent material.

Climate[link]

Climate is the dominate factor in soil formation, and soils show the distinctive characteristics of the climate zones in which they form.^[22] Mineral precipitation and temperature are the primary climate influences on soil formation.

The direct influence of climate include^[23]:

1. A shallow accumulation of lime in low rainfall areas as caliche.
2. Formation of acid soils in humid areas.
3. Erosion of soils on steep hillsides.
4. Deposition of eroded materials downstream
5. Very intense chemical weathering, leaching, and erosion in warm and humid regions where soil does not freeze.

Climate directly affects the rate of weathering and leaching. Soil is said to be formed when detectable layers of clays, organic colloids, carbonates, or soluble salts have been moved downward. Wind moves sand and smaller particles, especially in arid regions where there is little plant cover. The type and amount of precipitation influence soil formation by affecting the movement of ions and particles through the soil, and aid in the development of different soil profiles. Soil profiles are more distinct in wet and cool climates, where organic materials may accumulate, than those in wet, warm climates where organic materials are rapidly consumed. The effectiveness of water in weathering parent rock material depends on seasonal and daily temperature fluctuations. Cycles of freezing and thawing constitute an effective mechanism that breaks up rocks and other consolidated materials.

Climate indirectly influences soil formation by the effect of vegetation cover, biological activity, hence the rates of chemical reactions in the soil.

Topography[link]

The topography or relief characterized by the inclination of the surface determines the rate of precipitation runoff and rate of formation and erosion of the surface soil profiles. Steep slopes allow rapid runoff and erosion of the top soil profiles and little mineral deposition in lower profiles. Depressions allow the accumulation of water, minerals and organic matter and in the extreme, the resulting soils will be saline marshes or peat bogs. Intermediate topography affords the best conditions for the formation of an agriculturally productive soil.

Organisms[link]

Plants, animals, fungi, bacteria and humans affect soil formation (see soil biomantle and stonelayer). Animals and micro-organisms mix soils as they form burrows and pores, allowing moisture and gases to move about. In the same way, plant roots open channels in soils. Plants with deep taproots can penetrate many meters through the different soil layers to bring up nutrients from deeper in the profile. Plants with fibrous roots that spread out near the soil surface have roots that are easily decomposed, adding organic matter. Micro-organisms, including fungi and bacteria, effect chemical exchanges between roots and soil and act as a reserve of nutrients. Humans can impact soil formation by removing vegetation cover with erosion as the result. They can also mix the different soil layers, restarting the soil formation process as less weathered material is mixed with the more developed upper layers. Some soils may contain up to one million species of microbes per gram (most of those species being unknown), making soil the most abundant ecosystem on Earth.^[24]

Vegetation impacts soils in numerous ways. It can prevent erosion caused by excessive rain and the resulting surface runoff. Plants shade soils, keeping them cooler and slowing evaporation of soil moisture, or conversely, by way of transpiration, plants can cause soils to lose moisture. Plants can form new chemicals that can break down minerals and improve soil structure. The type and amount of vegetation depends on climate, land form topography, soil characteristics, and biological factors. Soil factors such as density, depth, chemistry, pH, temperature and moisture greatly affect the type of plants that can grow in a given location. Dead plants and fallen leaves and stems begin their decomposition on the surface. There, organisms feed on them and mix the organic material with the upper soil layers; these added organic compounds become part of the soil formation process.

Time[link]

Time is a factor in the interactions of all the above. Over time, soils evolve features dependent on the other forming factors. Soil formation is a time-responsive process that is dependent on how the other factors interplay with each other. Soil is always changing. It takes about 800 to 1000 years for a 2.5 cm thick layer of fertile soil to be formed in nature. For example, recently deposited material from a flood exhibits no soil development because there has not been enough time for the material to form a structure that further defines soil. The original soil surface is buried, and the formation process must begin anew for this deposit. Over a period of time from hundreds to thousands of years the soil will develop a profile that depends on the intensities of biota and climate. While soil can achieve relative stability of its properties for extended periods, the soil life cycle ultimately ends in soil conditions that leave it vulnerable to erosion. Despite the inevitability of soil retrogression and degradation, most soil cycles are long.

Soil-forming factors continue to affect soils during their existence, even on “stable” landscapes that are long-enduring, some for millions of years. Materials are deposited on top and materials are blown or washed from the surface. With additions, removals and alterations, soils are always subject to new conditions. Whether these are slow or rapid changes depend on climate, topography and biological activity.

Physical properties of soils[link]

The physical properties of soils, in their order of decreasing importance, are its texture, structure, density, porosity, consistency, temperature, color and resistivity. These determine the availability of oxygen in the soil and ability of water to infiltrate and be held in the soil. Soil texture is determined by the relative proportion of the three kinds of soil particles, called soil "separates": sand, silt and clay. Larger soil structures are created from the separates when iron oxides, carbonates, clay and silica with the organic constituent humus, coat particles and cause them to adhere into relatively stable secondary structures called "peds". Soil density, particularly bulk density, is a measure of the soil compaction. Soil porosity consists of the part of the volume occupied by air and water. Consistency is the ability of soil to stick together. Soil temperature and color are self defining. Resistivity refers to the resistance to conduction of electric currents and affects the rate of corrosion of metal structures and concrete. Soil properties may change through the depth of a particular soil profile with each identifiable layer in the profile.

Texture[link]

Soil types by clay, silt and sand composition as used by the USDA

Iron rich soil near Paint Pots in Kootenay National Park of Canada.

The mineral components of soil, sand, silt and clay determine a soils texture. In the illustrated textural classification triangle the only soil that does not exhibit one of those predominately is called "loam." While even pure sand, silt or clay may be considered a soil, from the perspective of food production a loam soil with a small amount of organic material is considered ideal. The mineral constituents of a loam soil might be 40% sand, 40% silt and the balance 20% clay by weight. Soil texture affects soil behavior, in particular its retention capacity for nutrients and water.^[25]

Sand and silt are the products of physical and chemical weathering, while clay is frequently the precipitated product of chemical weathering. Clay on the other hand is a product of chemical weathering and often forms as a secondary mineral from dissolved minerals that precipitate from solution. It is the specific surface area of soil particles and the unbalanced ionic charges within them that determine their role in the cation exchange potential of soil, hence its fertility. Sand is least active followed by silt; clay is the most active. Sand has its greatest benefit to soil by resisting compaction and increasing porosity. Silt, with its higher specific surface area, is more chemically active than sand, but the clay content, with its very high specific surface area and generally large number of negative charges, gives clay its great retention capacity for nutrients and water. Clay soils resist wind and water erosion better than silty and sandy soils, as the particles are bonded to each other.

Sand is the most stable of the mineral components of soil; it consists of rock fragments, primarily quartz particles, ranging in size from 2.0 mm to 0.05 mm. Sand is largely inert but plays an important part in holding open soil. Silt ranges in size from 0.05 mm to 0.002 mm. Silt is mineralogically like sand but is more active than sand due to its larger surface area. Clay is the most important component of mineral soil due to its net negative charge and ability to hold cations. Clay cannot be resolved by optical microscopes; it ranges in size from 0.002 mm or less.^[26] In medium-textured soils, clay is often washed downward through the soil profile and accumulates in the subsoil.

Soil components larger than 2.0 mm are classed as rock and gravel and are removed before determining the percentages of the remaining components and the texture class of the soil but are included in the name. For example, a sandy loam soil with 20% gravel would be called gravely sandy loam.

When the organic component of a soil is substantial, the soil is called organic soil rather than mineral soil. A soil is called organic if:

1. Mineral fraction is 0% clay and organic matter is 20% or more.
2. Mineral fraction is 0% to 50% clay and organic matter is between 20% to 30%.
3. Mineral fraction is 50% or more clay and organic matter 30% or more.^[27]

Structure[link]

The clumping of the soil textural components of sand, silt and clay forms aggregates and the further association of those aggregates into larger units forms soil structures called peds. The adhesion of those soil components by organic substances, iron oxides, carbonates, clays and silica, and by the breakage of those aggregates due to expansion-contraction, freezing-thawing, and wetting-drying cycles forms soil into distinct geometric forms. These peds evolve into units that may have various shapes, sizes and degrees of development.^[28] A soil clod is not a ped but rather a mass of soil that results from mechanical disturbance. The soil structure affects aeration, water movement, conduction of heat, resistance to erosion and plant root growth. Water has the strongest effect on soil structure due to its solution and precipitation of minerals and its effect on plant growth.

Soil structure often gives clues to its texture, organic matter content, biological activity, past soil evolution, human use, and chemical and mineralogical conditions under which the soil formed. While texture, is defined by the mineral component of a soil and is an innate property of the soil and does not change with agricultural activities, soil structure can be improved or destroyed by our choice and timing of farming practices.

Soil Structural Classes:^[29]

1. Types: Shape and arrangement of peds

a. Platy: Peds are flattened one atop the other 1-10 mm thick.

Found in the A-horizon of forest soils and lake sedimentation.

b. Prismatic and Columnar: Prismlike peds are long in the

vertical dimension, 10-100 mm wide. Prismatic peds have flat

tops, columnar peds have rounded tops. Tend to form in the B-

horizon in high sodium soil where clay has accumulated.

c. Angular and subangular: Blocky peds are imperfect cubes,

5-50 mm, angular have sharp edges, subangular have rounded

edges. Tend to form in the B-horizon where clay has

accumulated and indicate poor water penetration.

d. Granular and Crumb: Spheroid peds of polyhedrons, 1-10 mm,

often found in the A-horizon in the presence of organic

material. Crumb peds are more porous and are considered ideal.

2.Classes: Size of peds whose ranges depend upon the above type

a. Very fine or very thin: <1 mm platy and spherical; <5 mm

blocky; <10 mm prismlike.

b. Fine or thin: 1-2 mm platy, and spherical; 5-10 mm blocky;

10-20 mm prismlike.

c. Medium: 2-5 mm platy, granular; 10-20 mm blocky; 20-50

prismlike.

d. Coarse or thick: 5-10 mm platy, granular; 20-50 mm blocky;

50-100 mm prismlike.

e. Very coarse or very thick: >10 mm platy, granular; >50 mm

blocky; >100 mm prismlike.

3. Grades: Is a measure of the degree of development or cementation within the

peds that results in their strength and stability.

a. Weak: Weak cementation allows peds to fall apart into the

three constituents of sand, silt and clay.

b. Moderate: Peds are not distinct in undisturbed soil but when

removed they break into aggregates, some broken aggregates and

little unaggregated material. This is considered ideal.

c. Strong:Peds are distinct before removed from the profile and

do not break apart easily.

d. Structureless: Soil is entirely cemented together in one

great mass such as slabs of clay or no cementation at all such

as with sand.

At the largest scale, the forces that shape a soil's structure result from swelling and shrinkage that initially tend to act horizontally, causing vertically oriented prismatic peds. Clayey soil will induce horizontal cracks reducing columns to blocky peds. Roots, rodents, worms and freezing-thawing further break the peds into a spherical shape.

At a smaller scale, plant roots extend into voids and remove water and cause the open spaces to increase, and further decrease physical aggregation size. At the same time roots, fungal hyphea and earthworms create microscopic tunnels that break up peds.

At an even lower scale, soil aggregation continues as bacteria and fungi exude sticky polysaccharides that bind soil into small peds. The addition of the raw organic matter that bacteria and fungi feed upon encourages the formation of this desirable soil structure.

At the lowest scale, the soil chemistry affects the aggregation or dispersal of soil particles. The clay particles contain polyvalent cations that give faces of clay layers a net negative charge. At the same time the edges of the clay plates have a slight positive charge, thereby allowing the edges to adhere to the faces of other clay particles or to flocculate. On the other hand, when monovalent ions such as sodium invade and displace the polyvalent cations they weaken the positive charges on the edges, while the negative surface charges are relatively strengthened. This leaves a net negative charge on the clay, causing them to push apart, and so prevents the flocculation of clay particles into larger assemblages. As a result, the clay disperses and settles into voids between peds causing them to close. In this way the soil aggregation is destroyed and is made impenetrable to air and water. Such sodic soil tends to form columnar structures near the surface.^[30]

Density[link]

Density is the weight per unit volume of an object. Particle density is the density of the mineral particles that make up a soil i.e. excluding pore space and organic material. Particle density averages approximately 2.65 g/cc (165 lbm/ft³). Soil bulk density, a dry weight, includes air space and organic materials of the soil volume. A high bulk density indicates either compaction of the soil or high sand content. The bulk density of cultivated loam is about 1.1 to 1.4 g/cc (for comparison water is 1.0 g/cc).^[31] A lower bulk density by itself does not indicate suitability for plant growth due to the influence of soil texture and structure.

Porosity[link]

Pore space is that part of the bulk volume not occupied by either mineral or organic matter but is open space occupied by either air or water. Ideally, the total pore space should be 50% of the soil volume. The air space is needed to supply oxygen to organisms decomposing organic matter, humus and plant roots. Pore space also allows the movement and storage of water and dissolved nutrients.

There are four categories of pores:

1. Very fine pores: < 2 microns
2. Fine pores: 2-20 microns
3. Medium pores: 20-200 microns
4. Coarse pores: 200 microns-0.2 mm

In comparison, the root hairs are 8 to 12 microns in diameter. When pore space is less than 30 microns, the forces of attraction that hold water in place are greater than those acting to drain the water. At that point, soil becomes water logged and it cannot breathe. For a growing plant, pore size is of greater importance than total pore space. A medium textured loam provides the ideal balance of pore sizes. Having large pore spaces that allow rapid air and water movement is superior to smaller pore space but has a greater percentage pore space.^[33] Tillage has the short term benefit of temporarily increasing the number of pores of largest size but in the end those will be degraded by the destruction of soil aggregation.^[34]

Consistency[link]

Consistency is the ability of soil to stick together and resist fragmentation. It is of use in predicting cultivation problems and engineering of foundations. Consistency is measured at three moisture conditions: air-dry, moist and wet. The measures of consistency border on subjective as they employ the "feel" of the soil in those states. A soil's resistance to fragmentation and crumbling is made in the dry state by rubbing the sample. Its resistance to shearing forces is made in the moist state by thumb and finger pressure. Finally, a soils plasticity is measured in the wet state by molding with the hand.

The terms used to describe soil in those three moisture states and a last state of no agricultural value are as follows:

1. Consistency of Dry Soil: loose, soft, hard, extremely hard.
2. Consistency of Moist Soil: loose, friable, firm, extremely firm.
3. Consistency of Wet Soil: non-sticky, sticky or non-plastic, plastic
4. Consistency of Cemented Soil: weakly cemented, indurated (cemented)

Soil consistency is useful in estimating the ability of soil to support buildings and roads. More precise measures of soil strength are often made prior to construction.

Temperature[link]

Soil temperature regulates germination, root growth and availability of nutrients. Soil temperatures range from permafrost at a few inches below the surface to 38 C (100 F) in Hawaii on a warm day. The color of the ground cover and insulating ability have a strong influence on soil temperature. Snow cover and heavy mulching will reflect light and slow the warming of the soil, but at the same time reduce the fluctuations in the surface temperature.

Below 50 cm (20 in), soil temperature seldom changes and can be approximated by adding 1.8 C (2 F) degrees to the mean annual air temperature

Most often, soil temperatures must be accepted and agricultural activities adapted to them to:

1. maximize germination and growth by timing of planting.
2. optimize use of anhydrous ammonia by applying to soil below 10 C (50 F).
3. prevent heaving and thawing of frosts from damaging shallow rooted crops.
4. prevent damage to soil tilth by freezing of saturated soils.
5. improve uptake of phosphorus by plants.

Otherwise soil temperatures can be raised by drying soils or using clear plastic mulches. Organic mulches slow the warming of the soil.

Color[link]

Soil color is often the first impression one has when viewing soil. Striking colors and contrasting patterns are especially noticeable. The Red River (Mississippi watershed) carries sediment eroded from extensive reddish soils like Port Silt Loam in Oklahoma. The Yellow River in China carries yellow sediment from eroding loess soils. Mollisols in the Great Plains are darkened and enriched by organic matter. Podsols in boreal forests have highly contrasting layers due to acidity and leaching.

In general, color is determined by organic matter content, drainage conditions, and the degree of oxidation. Soil color, while easily discerned, has little use in predicting soil characteristics.^[35] It is of use in distinguishing boundaries within a soil profile, the origin of a soil's parent material, as an indication of wetness and waterlogged conditions, and as a qualitative means of measuring organic, salt and carbonate contents of soils. Color is recorded in the Munsell color system as for instance 10YR3/4.

Soil color is primarily influenced by soil mineralogy. Many soil colors are due to various iron minerals. The development and distribution of color in a soil profile result from chemical and biological weathering, especially redox reactions. As the primary minerals in soil parent material weather, the elements combine into new and colorful compounds. Iron forms secondary minerals with a yellow or red color, organic matter decomposes into black and brown compounds, and manganese, sulfur and nitrogen can form black mineral deposits. These pigments can produce various color patterns within a soil. Aerobic conditions produce uniform or gradual color changes, while reducing environments (anaerobic) result in disrupted color flow with complex, mottled patterns and points of color concentration.^[36]

Resistivity[link]

Soil resistivity is a measure of a soil's ability to retard the conduction of an electric current. The electrical resistivity of soil can affect the rate of galvanic corrosion of metallic structures in contact with the soil. Higher moisture content or increased electrolyte concentration can lower the resistivity and increase the conductivity thereby increasing the rate of corrosion.^[37][38] Soil resistivity values typically range from about 2 to 1000 Ω·m, but more extreme values are not unusual.^[39]

Soil water[link]

Water effects soil formation, structure, stability and erosion but is of primary concern with respect to plant growth. Water is essential to plants for four reasons:

1. It constitutes 85%-95% of the plants protoplasm.
2. It is essential for photosynthesis.
3. It is the solvent in which nutrients are carried to, into and throughout the plant.
4. It provides the turgidity by which the plant keeps itself in proper position.^[40]

In addition, water alters the soil profile by dissolving and redepositing minerals, often at lower levels, and possibly leaving the soil sterile in the case of extreme rainfall and drainage. In a loam soil, solids constitute half the volume, air one-quarter of the volume, and water one-quarter of the volume of which only half of that water will be available to most plants.

Water retention forces[link]

Water is retained in a soil when the adhesive force of attraction of water for soil particles and the cohesive forces water feels for itself are capable of resisting the force of gravity that tends to drain water from the soil. When a field is flooded, the air space is displaced by water. The field will drain under the force of gravity until it reaches what is called field capacity at which point the smallest pores are filled with water and the largest with water and air.^[41] The total amount of water held when field capacity is reached is a function of the specific surface area of the soil particles. As a result, high clay and high organic soils have higher field capacities. The total force required to pull, or push water out of soil is given the term suction and usually expressed in units of bars (10⁵ pascal) which is just a little less than one-atmosphere pressure. Alternatively, the terms tension or moisture potential may be used.^[42]

Moisture classification[link]

The forces with which water is held in soils determines its availability to plants. Forces of adhesion hold water strongly to mineral and humus surfaces and less strongly to itself by cohesive forces. A plant's root may penetrate a very small volume of water that is adhering to soil and be able initially to draw water in that is only lightly held by the cohesive forces . But as the droplet is drawn down, the forces of adhesion of the water for the soil particles make reducing the volume of water increasingly difficult until the plant cannot produce sufficient suction to use the remaining water. The remaining water is considered unavailable. The amount of available water depends upon the soil texture and humus amounts and the type of plant. Cacti can for example, produce greater suction than can agricultural crop plants.

The following description applies to a loam soil and agricultural crops. When a field is flooded it is called saturated and all available air space is occupied by water. The suction required to draw water into a plant root is zero. As the field drains under the influence of gravity (drained water is called gravitational water or drain-able water), the suction required to be produced by the plant to use such water increases to 1/3 bar. At that point, the soil is said to have reached field capacity, and plants that use the water must produce increasingly higher suction, finally up to 15 bar. At 15 bar suction the soil water amount is called wilting percent. At that suction the plant cannot sustain its water needs as water is still being lost from the plant by transpiration; the plant's turgidity is lost, and it wilts. The next level, called air-dry, occurs at 1000 bar suction. Finally the oven dry condition is reached and at 10,000 bar suction. All water below wilting percentage is called unavailable water.^[43]

Soil moisture content[link]

The amount of water remaining in a soil drained to field capacity and the amount that is available is a function of the soil type. Sandy soil will retain very little water while clay will hold the maximum amount. The time required to drain a field from flooded condition for a clay loam that begins at 43% water by weight to a field capacity of 21.5% is six days whereas for a sand loam that is flooded to its maximum of 22% water, will take two days to reach field capacity of 11.3% water. The available water for the clay loam might be 11.3% whereas for the sand loam it might be only 7.9% by weight.^[44]

The above are average values for the soil textures as the percentage of sand, silt, and clay vary within the listed soil textures.

Water flow in soils[link]

Water moves through soil due to the force of gravity, osmosis and capillarity. At zero bar suction to one-third bar suction, water moves through soil due to gravity and is called saturated flow. At higher suction, water movement is called unsaturated flow.^[46]

Water infiltration into soil is controlled by six factors:

1. Soil texture
2. Soil structure. Fine-textured soil with granular structure are most favorable.
3. The amount of organic matter. Coarse matter is best and if on the surface helps prevent the destruction of soil structure and the creation of crusts.
4. Depth of soil to impervious layers such as hardpans or bedrock.
5. The amount of water already in the soil.
6. Soil temperature. Warm soils take in water faster while frozen soils may not be able to absorb depending on the type of freezing.^[47]

Water infiltration rates range from 0.25 cm per hour for high clay soils to 2.5 cm per hour for sand, and well stabilized and aggregated soil structures.^[48]

Saturated flow[link]

Once soil is completely wetted, any more water will move downward, or percolate, carrying with it clay, humus and nutrients, primarily cations, out of the range of plant roots and result in acid soil conditions. In order of decreasing solubility, the leached nutrients are:

• Calcium
• Magnesium, Sulfur, Potassium; depending upon soil composition.
• Nitrogen; usually little, unless nitrate fertilizer was applied recently.
• Phosphorus; very little as its forms in soil are of low solubility.^[49]

In the United States percolation water due to rainfall ranges from zero inches just east of the Rocky Mountains to twenty or more inches in the Appalachian Mountains and the north coast of the Gulf of Mexico.^[45]

Unsaturated flow[link]

At suctions less than one-third bar, water moves in all directions in unsaturated flow at a rate that is dependent on the square of the diameter of the water filled pores. Water is pushed by pressure gradients, from the point of its application where it is saturated locally, and pulled by capillary action due to adhesion force of water for the soil solids, producing a suction gradient from wet toward drier soil. Doubling the diameter of the pores increases the flow rate by a factor of four. Large pores drained by gravity and not filled with water do not greatly increase the flow rate for unsaturated flow. Water flow is primarily from coarse textured soil into fine-textured soil and moves most slowly through fine-textured soils such a clay.^[50]

Water uptake by plants[link]

Of equal importance to the storage and movement of water in soil is the means by which plants acquire it and their nutrients. Ninety percent of water is taken up by plants as passive absorption caused by the pulling force of water evaporating (transpiring) from the long column of water that leads from its roots to its leaves. In addition, the high concentration of salts within the plant roots create an osmotic pressure gradient that pushes soil water into the roots. Osmotic absorption becomes more important during times of low water transpiration at night (lower temperatures) or due to high humidity during the day. It is the process which causes guttation.^[51]

Root extension is vital for plant survival. A study of a single winter rye plant grown for four months in one cubic foot of loam soil showed that the plant developed 13,800,000 roots of 385 miles and 2,550 square feet of surface area and 14 billion hair roots of 6,600 miles and 4,320 square feet of root area, for a total surface area of 6,870 square feet. The total surface area of the loam soil was estimated to be 560,000 square feet.^[52] In other words the roots were in contact with only 1.2% of the soil. Roots must seek out water as the unsaturated flow of water in soil can move only at a rate of up to 2.5 cm per day, as a result they are constantly dying and growing, as they seek out high concentrations of soil moisture.

Insufficient soil moisture to the point of wilting will cause permanent damage and crop yields will suffer. When grain sorghum was exposed to soil suction as low as 13.0 bar during the seed head emergence through bloom and seed set stages of growth, the production was reduced by 34%.^[53]

Consumptive use and water efficiency[link]

Only a small fraction (0.1% to 1%) of the water used by a plant is held within the plant. Transpiration of water from the plant is the majority of water's use, while evaporation from the soil surface is also substantial. Transpiration plus evaporative soil moisture loss is called evapotranspiration. Evapotranspiration, plus water held in the plant totals to consumptive use which is nearly identical to evapotranspiration.^[53]

The total water used in an agricultural field includes runoff, drainage, and consumptive use. The use of loose mulches will reduce evaporative losses for a period after a field is irrigated but in the end the total evaporative loss will approach that of an uncovered soil. The benefit from mulch is to keep the moisture available during the seedling stage. Water use efficiency is measured by transpiration ratio which is the ratio of the total water transpired by a plant to the dry weight of the harvested plant at a particular locale. Alfalfa may have a transpiration ratio of 500 (for a particular location) and as a result 500 kilograms of water will produce one kilogram of dry alfalfa. Transpiration ratios for crops range from 300 to 700.^[54]

Soil Atmosphere[link]

The atmosphere of soil is radically different from that of the atmosphere above. The consumption of oxygen by microbes and plant roots and their release of carbon dioxide decreases oxygen, and increases carbon dioxide concentration. Atmospheric CO₂ concentration is 0.03% but in the soil pore space it may range from 10 to 100 times that level. In addition the void is saturated with water vapor. Adequate porosity is necessary not just to allow the penetration of water but also to allow gasses to diffuse in and out. Movement of gasses is by diffusion from high concentrations to lower. Oxygen diffuses in and is consumed and excess levels of carbon dioxide, which can become toxic, diffuse out with other gasses as well as water. Soil texture and its structure strongly affects soil porosity and gas diffusion.^[55] Platy and compacted soils impede gas flow and a deficiency of oxygen may encourage anaerobic bacteria to reduce nitrate to N₂, N₂O, and NO, which is then lost to the atmosphere. Aerated soil is also a net sink of methane CH₄ but a net producer of greenhouse gases when soils are depleted of oxygen and subject to elevated temperatures.^[56]

Chemical and colloidal properties[link]

The chemistry of soil determines the availability of nutrients, the health of microbial populations, and its physical properties. Soil chemistry determines its corrosivity, stability, and ability to absorb pollutants and to filter water. It is the surface chemistry of clays and humus colloids that determines soil's chemical properties. The very high specific surface area of colloids gives soil its great ability to hold and release cations in what is referred to as cation exchange. Cation exchange capacity is the amount of exchangeable cations per unit weight of dry soil and is expressed in terms of milliequivalents of hydrogen ion per 100 grams of soil. “A colloid is a small, insoluble, nondiffusible particle larger than a molecule but small enough to remain suspended in a fluid medium without settling. Most soils contain organic colloidal particles as well as the inorganic colloidal particles of clays.”^[57]

Soil clays[link]

Due to its high specific surface area, clay is the most active mineral component of soil. It is a colloidal and crystalline material. In soils, clay is defined in a physical sense as any mineral particle less that two microns in effective diameter. Chemically, clay is a range of minerals with certain reactive properties. Clay is also a soil textural class. Many soil minerals, such a gypsum, carbonates or quartz, are small enough to be classified physically as clay but they do not afford the same utility chemically as do clay minerals.^[58]

Clay was once thought to be very small particles of quartz, feldspar, mica, hornblende or augite, but it is now known to be (with the exception of mica based clays) a precipitate with a mineralogical composition different from its parent materials and is classed as a secondary mineral. The type of clay that is formed is a function of the parent material and the composition of the minerals in solution. Mica based clays result from a modification of the primary mica mineral in such a way that it behaves and is classed as a clay. Most clays are crystalline but some are amorphous. The clays of soil are a mixture of the various types of clay but one type predominates.

Most clays are crystalline and most are made up of three or four planes of oxygen held together by planes of aluminum and silicon by way of ionic bonds that together form a single layer of clay. It is the spacial arrangement of the oxygen atoms that determines clay's structure. Half of the weight of clay is oxygen but on a volume basis oxygen is ninety percent.^[59] The layers of clay are sometimes held together through hydrogen bonds or potassium bridges and as a result are less swelling in the presence of water. Other clays layers are loosely attached and will swell greatly when water intervenes between the layers.

There are three groups of clays:

1. Crystalline alumino-silica clays, montmorillonite, illite, vermiculite, chlorite, kaolinite.
2. Amorphous clays, are young mixtures of silica (SiO₂-OH) and alumina (Al(OH)₃) but have not had time to form regular crystals.
3. Sesquioxide clays are highly leached clays of iron, aluminum and titanium oxides.^[60]

Silica clays[link]

Alumino-silica clays are characterized by their regular crystalline structure. Oxygen in ionic bonds with silicon forms a tetrahedral coordination that in turn forms sheets of silica. Two sheets of silica are bonded together by a plane of aluminum that form an octahedral coordination, called alumina, with the oxygens of the silica sheet above and that below it. Hydroxyl ions (OH^-) sometimes substitute for oxygen. As much as one fourth of the aluminum Al³⁺ may be substituted by Zn²⁺, Mg²⁺ or Fe²⁺ and Si⁴⁺ may be substituted by Al³⁺. The substitution of lower valence ions for higher valence ions (isomorphic substitution) gives clay a net negative charge that attracts and holds cations some of which are of value for plant growth.

• Montmorillonite clay is made of four planes of oxygen with two silicon and one central aluminum planes intervening. The alumino-silicate montmorillonite clay are said to have a 2:1 ratio of silicon to aluminum. The seven planes together form a single layer of montmorillonite. The layers are weakly held and water may intervene, causing the clay to swell up to ten times its dry volume. It occurs in soils that have had little leaching. Montmorillonite is found in arid regions. The entire surface is exposed and available for surface reactions and has high cation exchange capacity.

• Illite is a 2:1 clay similar in structure to montmorillonite but has potassium bridges between the clay layers and the degree of swelling depends on the degree of weathering of the potassium. The active surface area is reduced due to the potassium bonds. Illite originates from the modification of mica, a primary mineral. It is often found together with montmorillonite with its primary minerals. It has moderate cation exchange capacity.

• Vermiculite is a mica based clay similar to illite but the layers of clay are held more loosely together by hydrated magnesium and will swell, but not a much as montmorillonite. It has very high cation exchange capacity.

• Chlorite is similar to vermiculite but the loose bonding by occasional hydrated magnesium is replaced by a hydrated magnesium sheet, firmly bonding the planes above and below it. It has two planes of silicon, one of aluminum and one of magnesium; hence it is a 2:2 clay. Chlorite does not swell and it has low cation exchange capacity.

• Kaolinite is very common; more common than montmorillonite in acid soils. It has one silica and one alumina sheet per layer; hence it is a 1:1 type clay. One layer of oxygen is replace with hydroxyls which produces strong hydrogen bonds to the oxygen in the next layer of clay. As a result kaolinite does not swell in water and has a low surface area, and as almost no isomorphic substitution has occurred it has a low cation exchange capacity. Where rainfall is high, acid soils selectively leaches more silica than it does alumina from the original clays leaving kaolinite. Even heavier weathering results in sesquioxide clays.

Amorphous clays[link]

Amorphous clays are common in volcanic ash. They are mixtures of alumina and silica that have not formed the ordered crystal shape of alumino-silica clays that time would provide. The majority of their negative charges originates from hydroxyl ions which can gain or lose a hydrogen ion (H⁺) hence buffer soil pH. They may have either a negative charge provided by the attached hydroxyl ion (OH^-), that can attract a cation, or lose the hydrogen of the hydroxyl to solution and display a positive charge that can attract anions. As a result they may display either high cation exchange capacity, in an acid soil solution, or high anion exchange capacity, in a basic soil solution.

Sesquioxide clays[link]

Sesquioxide clays are a product of heavy rainfall, that has leached most of the silica and alumina from alumino-silica clay, leaving the less soluble oxides of iron Fe₂O₃ and iron hydroxide (Fe(OH)₃) and aluminum hydroxides (Al(OH)₃). Sesqi is Latin meaning one and one-half; there are three parts oxygen to two parts iron or aluminum, hence the ratio is one and one-half. They are hydrated and act as either amorphous or crystalline. It takes hundreds of thousands of years of leaching to create sesquioxide clays. They are not sticky and do not swell and soils high in them behave much like sand and can rapidly absorb water. They are able to hold large quantities of phosphates. Sesquioxides have low cation exchange capacity.^[61][62]

Organic colloids[link]

Humus is the penultimate state of decomposition of organic matter; while it may linger for a thousand years, on the larger scale of the age of the other soil components, it is temporary. It is composed of the very stable lignins (30%) and complex sugars (polyuronides, 30%). On a dry weight basis, the cation exchange capacity of humus is many times greater than that of clay. Plant roots also have cation exchange sites.

Cation and anion exchange[link]

Cation exchange, between colloids and soil water, buffers (moderates) soil pH, alters soil structure, and purifies percolating water by adsorbing cations of all types, both useful and harmful.

The negative charges on a colloid particle makes it able to hold cations to its surface. The charges result from four sources.

1. Isomorphous substitution occurs in clay when lower valence cations substitute for higher valence cations in the crystal structure. Substitutions in the outermost layers is more effective than for the innermost layers as the charge strength drops off as the square of the distance. The net result is a negative charge.
2. Edge-of-clay oxygen atoms are not in balance ionically as the tetrahedral and octahedral structures are incomplete at the edges of clay.
3. Hydrogens of the clay hydroxyls may be ionized into solution leaving an oxygen with a negative charge.
4. Hydrogens of humus hydroxyl groups may be ionized into solution leaving an oxygen with a negative charge.^[63]

Cations held to the negatively charged colloids resist being washed downward by water and out of reach of plants roots, thereby saving the fertility of soils in areas of moderate rainfall and low temperatures.

There is a hierarchy in the process of cation exchange on colloids as they differ in the strength of adsorption and their ability to replace one another. If present in equal amounts:

Al³⁺ replaces H⁺ replaces Ca²⁺ replaces Mg²⁺ replaces K⁺ same as NH⁴⁺ replaces Na^+[57]

If one cation is added in large amounts it may replace another by the sheer force of it numbers (mass action). This is largely what occurs with the addition of fertilizer.

As the soil solution becomes more acidic, the other cations bound to colloids are pushed into solution. This is caused by the ionization of hydroxyl groups on the surface of soil colloids in what is describes as pH dependent charges. Unlike permanent charges developed by isomorphous substitution, pH-dependent charges are variable and increase with increasing pH.^[64] As a result those cations can be made available to plants but also able to be leached from the soil, possibly making the soil less fertile. Plants will excrete H⁺ to the soil and to replace cations on the colloids, making the cations available to the plant.

Cation exchange capacity (CEC)[link]

Cation exchange capacity should be thought of as the soils ability to remove cations from the soil water solution and sequester those to be exchanged later as the plant roots release hydrogen ions to the solution. CEC is the amount of exchangeable hydrogen cation (H⁺) that will combine with 100 grams dry weight of soil and whose measure is one milliequivalents per 100 grams of soil (1 meq/100 g). Hydrogen ions have a single charge and one-thousandth of a gram of hydrogen ions per 100 grams dry soil gives a measure of one milliequivalent of hydrogen ion. Calcium with an atomic weight 40 times that of hydrogen and with a valence of two, converts to (40/2) x 1 milliequivalent = 20 milliequivalents of hydrogen ion per 100 grams of dry soil or 20 meq/100 g.^[65] The modern measure of CEC is expressed as centimoles of positive charge per kilogram (cmol/kg), of oven dry soil.

Most of the soil's CEC occurs on clay and humus colloids and the lack of those in hot, humid, wet climates, due to leaching and decomposition respectively, explains the sterility of tropical soils.

Anion exchange capacity (AEC)[link]

Anion exchange capacity should be thought of as the soils ability to remove anions from the soil water solution and sequester those for later exchange as the plant roots release carbonate anions to the solution. Those colloids that have low CEC tend to have some AEC. Amorphous and sesquioxide clays have the highest AEC followed by the iron oxides. The levels of AEC is much lower than for CEC. Phosphates tend to be held at anion exchange sites.

Iron and aluminum hydroxide clays are able to exchange their hydroxide anions (OH^-) for other anions. The order reflecting the strength of anion adhesion is as follows:

H₂PO₄^- replaces SO₄^2- replaces NO₃^- replaces Cl^-

The amount of exchangeable anions are of a magnitude of tenths to a few milliequivalents per 100 g dry soil.^[67] As the pH rises there is more hydroxyls that will displace anions from the colloids and force them into solution and out of storage, hence the AEC decreases.

Soil reaction (pH)[link]

Soil reactivity is expressed in terms of pH and is a measure of the acidity and alkaninity of the soil. More precisely, it is a measure of hydrogen ion concentration in an aqueous solution and ranges in values from 0 to 14 (acidic to basic) but practically speaking for soils, pH ranges from 3.5 to 9.5 as pH values beyond those extremes are toxic to life forms.

Soil pH[link]

At 25°C an aqueous solution that has a pH of 3.5 has 10^-3.5 moles hydrogen ions per liter of solution (and also 10^-10.5 mole/liter OH^-) . A pH of 7, defined as neutral, has 10⁻⁷ moles hydrogen ions per liter of solution and also 10⁻⁷ moles of OH^- per liter; since the two concentrations are equal they are said to neutralize each other. A pH of 9.5 is 10^-9.5 moles hydrogen ions per liter of solution (and also 10^-3.5 mole per liter OH^-) . A pH of 3.5 has one million times more hydrogen ions per liter than a solution with pH of 9.5 (9.5 - 3.5 = 6 or 10⁶) and is more acidic.^[68]

Plants differ in their nutrient needs and the effect of pH is to remove from the soil or make available certain ions. High acid soils tend to have toxic amounts of aluminum and manganese. Plants that need calcium need moderate alkalinity but most minerals are more soluble in acid soils. Soil organisms are hindered by high acidity and most agricultural crops do best on mineral soils of pH 6.5 and organic soils of pH of 5.5.

In high rainfall areas, soils tend to acidity as the basic cations are leached away by rain allowing the soil colloids to become saturated with hydrogen ions from naturally acid rain leaving the soil sterile. The addition of any more hydrogen ions or aluminum hydroxyl cations drives the pH even lower as the soil is left with no buffering capacity. In low rainfall areas, unleached calcium pushes pH to 8.5 and with the addition of exchangeable sodium, soils may reach pH 10. Beyond a pH of 9 plant growth is reduced. High pH results in low micro-nutrient mobility, but water soluble-chelates of those can supply the deficit. Sodium can be reduced by the addition of gypsum (calcium-sulphate).

Base saturation percentage[link]

There are acid-forming cations (hydrogen and aluminum) and there are base-forming cations. The fraction of the base-forming cations that occupy positions on the soil colloids is called the base saturation percentage. If a soil has a CEC of 20 meq and 5 meq are aluminum and hydrogen cations (acid-forming) the remainder (20-5 = 15 meq) are assumed occupied by base-forming cations, then the percentage base saturation is 15/20 x 100% = 75 percent (the compliment 25% is assumed acid-forming cations). When the soil pH is 7 (neutral) base saturation is 100 percent and there are no hydrogen ions stored on the colloids. Base saturation is almost in direct proportion to pH and except for its use in calculating the amount of lime needed to neutralize an acid soil it is of little use.

Buffering of soils[link]

The resistance of soil to changes in pH and available cations from the addition of acid or basic material is a measure of the buffering capacity of a soil and increases as the CEC increases. Hence, pure sand has almost no buffering ability. Buffering occurs by cation exchange and neutralization.

The addition of highly basic aqueous ammonia will cause the ammonium to displace hydrogen ions from the colloids and the end product is colloidally fixed ammonium and water, but no permanent change in pH.

The addition of lime, CaCO₃, will displace hydrogen ions from the soil colloids, causing the fixation of calcium, the freeing of CO₂ and leaving water, with no permanent change in pH.

The addition of carbonic acid (resulting from water and CO₂) will displace calcium from colloids, thereby fixing hydrogen ions, evolving water and slightly alkaline (temporary increase in pH) highly soluble calcium bicarbonate, which will precipitate as lime (CaCO₃) and water at a lower level in the soil. With the result of no permanent change in pH.

The general principal is that an increase in a particular cation in the soil water solution will cause that cation to be fixed to colloids (buffered) and a decrease in solution of that cation will cause it to be withdrawn from the colloid and moved into solution (buffered). The degree of buffering is limited by the CEC of the soil; the greater the CEC the greater the buffering capacity of the soil.

Nutrients[link]

There are sixteen nutrients essential for plant growth and reproduction. They are carbon, hydrogen, oxygen, nitrogen, phosphorus, potassium, sulfur, calcium, magnesium, iron, boron, manganese, copper, zinc, molybdenum, and chlorine. Nearly all plant nutrients are taken up in ionic forms from the water part of the soil solution as cations or as anions. Plants release bicarbonate and hydorxyl (OH^-) anions or hydrogen cations in an effort to cause nutrient ions to be freed from sequestration on colloids and so force them into the soil solution. Nitrogen ions and cations are stored in soil organic material and are made available to the plant roots by that material's decomposition by micro-organisms.^[69]

The mechanism of nutrient uptake[link]

All the nutrients with the exception of carbon are taken up by the plant through its roots. All those taken through the roots, with the exception of hydrogen which is derived from water, are taken up in the form ions. Carbon, in the form of carbon dioxide, enters primarily through the stomata of the leaves and where the plant releases oxygen as a byproduct of photosynthesis. All the hydrogen utilized by the plant originates from soil water and results in the release of further oxygen. Plants may have their nutrient needs supplemented by spraying a water solution of nutrients on their leaves, but nutrients are typically received through the roots by:

1. Mass flow.
2. Diffusion.
3. Root interception.

The nutrient needs of a plant may be carried to the plant by the movement the soil solution of water in a what is called mass flow. The absorption of nutrients by the roots from the water with which it is in contact, causes the concentration of nutrients in that area to be depleted. Nutrients then diffuse from areas with higher concentration to lower concentration, thereby bringing more nutrients near the roots. Plants also send out roots constantly to seek new sources of nutrients in a process called root interception. Meanwhile older less effective roots die back. Water is lifted to the leaves where it is lost by transpiration and in the process, it brings with it soil nutrients. A corn plant will use one quart of water per day at the height of its growing season.^[71]

Plants move ions of out of their roots in proportion to the amount of nutrients they move in. Hydrogen H⁺ is exchanged for cations and carbonate, HCO₃^- and hydroxide OH^- anions are exchanged for nutrient anions. Plants derive most of their anion nutrients from decomposing organic matter, which hold 95 percent of the nitrogen, 5 to 60 percent of the phosphorus and 80 percent of the sulfur. As plant roots remove nutrients from the soil water solution, nutrients are added as ions move off of clay and humus, are added from decomposition of soil minerals, and released by the decomposition of soil organic matter. Where crops are produced, the nutrients must be augmented by fertilizer or added organic matter.^[73]

Carbon[link]

Plants obtain their carbon from atmospheric carbon dioxide. A plant's weight is forty-five percent carbon. Elementally, carbon is 50% of plant material. Plant residues have a carbon to nitrogen ratio (C/N) of 50:1. As the soil organic material is digested by arthropods, and micro-organisms the C/N decreases as the carbonaceous material is metabolized and carbon dioxide (CO₂) byproduct is released and finds its way to the atmosphere. The nitrogen, however, is sequestered in the bodies of the live matter. Normal CO₂ concentration in the atmosphere is 0.03% which is probably the factor limiting plant growth. In a field of corn on a still day during high light conditions of the growing season, the CO₂ concentration drops very low but under such conditions the crop could use up to 20 times the normal concentration. The respiration of CO₂ by soil micro-organisms decomposing soil organic matter, contribute an important amount of CO₂ to the photo-synthesizing plants. Within the soil CO₂ concentration is 10 to 100 times atmospheric but may rise to toxic levels if the soil porosity is low or impeded by a flooded condition.^[74]

Nitrogen[link]

Nitrogen is the most critical element attained by plants from the soil and is a bottleneck in plant growth.^[75] Plants can use the nitrogen as either the cation, ammonium, NH₄⁺, or the anion nitrate, NO₃^-. Nitrogen is seldom missing in the soil but is in the form of raw organic material and cannot be used directly. Some micro-organisms are able to metabolize the organic matter and release ammonium in a process called mineralization. Others take free ammonium and oxidize it to nitrate. Particular bacteria are capable of metabolizing N₂ into the form of nitrate, in a process called nitrogen fixation. Both ammonium and nitrate can be lost from the soil by incorporation into the microbes living cells and temporarily immobilized or sequestered. Nitrate may also be lost when bacteria metabolize it to the gasses N₂, N₂O and so escape to the atmosphere in a process called denitrification. Nitrogen may be leached from the soil if it is in the form of nitrate and lost as a ammonia due to a chemical reaction of ammonium with alkaline soil called volatilization. Nitrogen is added to soil by rainfall. Ammonium may also be sequestered in clay by fixation.^[76][77]

Nitrogen gains[link]

In a process called mineralization, certain bacteria feed on organic matter, releasing ammonia (NH³) (which may be reduced to ammonium NH₄⁺) and other nutrients. As long as the carbon to nitrogen ratio (C/N) in the soil is above 30:1 nitrogen will be in short supply and other bacteria will feed on the ammonium and incorporate its nitrogen into their cells. In that form the nitrogen is said to be immobilized. Later when such bacteria die they too are mineralized. If the C/N is less than 15, ammonia is freed to the soil where it may be used by plants or certain bacteria may oxidize it to nitrate in a process called nitrification. Bacteria may on average add 25 pounds nitrogen per acre, and in an unfertilized field, it is the most important source of usable nitrogen. In a soil with 5 percent organic matter perhaps 2 to 5 percent of that is released to the soil by such decomposition. It occurs fastest in warm, moist, well aerated soil. The mineralization of 3 percent of a soil that is 4 percent organic matter would release 120 pounds of nitrogen as ammonium per acre.^[78]

In symbiotic fixation, Rhizobium bacteria, are capable of conversion of N₂ to nitrate in the process of nitrogen fixation. They have a symbiotic relationship with host plants wherein they supply the host with nitrogen and the host provides the bacteria with nutrients and a safe environment. It is estimated that such symbiotic bacteria in the root nodules of legumes add 45 to 250 pounds of nitrogen per acre per year, which may be sufficient for the crop. Other free living nitrogen fixing bacteria and blue-green algae live independently in the soil and release nitrate, when their dead bodies are converted by way of mineralization.^[79]

Some amount of usable nitrogen is fixed by lightning as nitric acid (HNO₃). Ammonia, NH₃, previously released from the soil or from combustion, and the nitric acid fall with precipitation in giving a total of amount of about five pounds nitrogen pounds per acre per year.^[80]

Nitrogen sequestration[link]

When bacteria feed on soluble forms of nitrogen (ammonium and nitrite) they temporarily sequester that nitrogen in their own bodies in a process called immobilization. At a later time when those bacteria die, their nitrogen may be released as ammonium by the processes of mineralization.

Protein material is easily broken down but the rate of its decomposition is slowed by its attachment to the crystalline structure of clay and held between the clay layers. The layers are small enough that bacteria cannot enter. Some organisms can exude extracellular enzymes that can act on the sequestered proteins. However, those enzymes too may be trapped on the clay crystals.

Ammonium fixation occurs when ammonium replaces the potassium ions that normally exist between the layers of clay such as illite or montmorillonite. Only a small fraction of nitrogen is held this way^[81]

Nitrogen losses[link]

Usable nitrogen may be lost from soils when in the form of nitrate, as it is easily leached. Further losses of nitrogen occur by denitrification, the process whereby soil bacteria convert nitrate, to nitrogen gas N₂ or N₂O. This occurs when poor soil aeration limits free oxygen, forcing bacteria to use the oxygen in nitrate for its respiratory process. Denitrification is increased when oxidizable organic material is available and when soils are warm and slightly acidic. Denitrification may vary throughout a soil as the aeration varies from place to place. The conversion of nitrate to gasses causes their loss to the atmosphere. Denitrification may cause the loss of 10 to 20 percent of the available nitrates within a day when conditions are favorable and the losses of up to 60 percent of applied nitrate as fertilizer.^[82]

Ammonium volatilization occurs when ammonium reacts chemically with an alkaline soil, converting NH₄⁺ to NH₃. The application of ammonium fertilizer to such a field can result in volatilization losses as much as 30 percent.^[83]

Phosphorus[link]

Phosphorus is the second most critical plant nutrient. The soil mineral apatite is the most common mineral source of phosphorus. While there is on average 1000 pounds of phosphorus per acre in the soil, it is generally in unavailable forms. The available portion of phosphorus is low as it is in the form of phosphates of low solubility. Total phosphorus is about 0.1 percent by weight of the soil but only one percent of that is available. Of the part available more than half comes from the mineralization of organic matter. Agricultural fields may need to be fertilized to make up for the phosphorus that has been removed in the crop.^[84]

When phosphorus does form solubilized ions of H₂PO₄^-, they rapidly form insoluble phosphates of calcium or hydrous oxides of iron and aluminum. Phosphorus is largely immobile in the soil and is not leached but actually builds up in the surface layer if not cropped. The application of soluble fertilizers to soils may result in zinc deficiencies as zinc phosphates form. Conversely, the application of zinc to soils may immobilize phosphorus as zinc phosphate. Lack of phosphorus may interfere with the normal opening of the plant leaf stomata resulting in plant temperatures 10 percent higher than normal. Phosphorus is most available when soil pH is 6.5 in mineral soils and 5.5 in organic soils.^[83]

Potassium[link]

The amount of potassium in a soil May 80,000 lb per acre of which only 150 lbs or 2 percent is available for plant growth. When solubilized, half will be held as exchangeable cations on clay while the other half is in the soil water solution. Potassium fixation occurs when soils dry and the potassium is bonded between layers of clay. Under certain conditions, dependent on the soil texture, intensity of drying, and initial amount of exchangeable potassium, the fixed percentage may be as much as 90 percent within ten minutes. In soils low in clay, potassium may be leached.^[85]

Calcium[link]

Calcium is 1 percent by weight of soils and is generally available but may be low as it is soluble and can be leached. It is generally available except in sandy and heavily leached soil or strongly acidic mineral soil. Calcium is supplied to the plant in the form of exchangeable ions and moderately soluble minerals. Calcium is more available on the soil colloids than is potassium because the common mineral, calcite CaCO₃ is more soluble than potassium bearing minerals.^[86]

Magnesium[link]

Magnesium is central to chlorophyll and aids in the uptake of phosphorus. The amount of magnesium beyond the minimum amount for plant health is not sufficient for forage animals. Magnesium is generally available, but is missing from some soils along the Gulf and Atlantic coasts of the United States.^[87]

Sulfur[link]

Most sulfur is made available to plants, like phosphorus, by its release from decomposing organic matter.^[87] Deficiencies may exist in some soils and if cropped sulfur need be added. A 15-ton crop of onions uses up to 19 pounds of sulfur and 4 tons of alfalfa uses 15 pounds per acre. Sulfur abundance varies with depth. In a sample of soils in Ohio, United States, the sulfur abundance varied with depths, 0-6 inches, 6-12 inches, 12-18 inches, 18-24 inches in the amounts: 1056, 830, 686, 528 pounds per acre respectively.

Micronutrients[link]

Micronutrients iron, manganese, zinc, copper, boron, chlorine, and molybdenum, refers to the plant needs not their abundance in soil, are required in very small amounts but are essential to plant health. They are generally available in the mineral component of the soil but the heavy application of phosphates can cause a deficiency in zinc and iron, by the formation of insoluble phosphates. Iron deficiency may result from excessive amounts of heavy metals or calcuium minerals (lime) in the soil. Excess amounts of soluble boron, molybdenum, and chloride are toxic.^[88]

Organic matter[link]

The organic soil matter includes all the dead plant material and all creatures live and dead. The living component of an acre of soil may contain 900 pounds of earthworms, 2400 pounds of fungi, 1500 pounds of bacteria, 133 pounds of protozoa, 890 pounds of arthropods and algae.^[89]

Most living things in soils, including plants, insects, bacteria and fungi, are dependent on organic matter for nutrients and energy. Soils have varying organic compounds in varying degrees of decomposition. Organic matter holds soils open, allowing the infiltration air and water and may hold as much twice its weight in water. Many soils, including desert and rocky-gravel soils, have no or little organic matter. Soils that are all organic matter, such as peat (histosols), are infertile.^[90] In its earliest stage of decomposition the original organic material is often called raw organic matter. The final stage of decomposition is called humus.

Humus[link]

Humus refers to organic matter that has been decomposed by bacteria, fungi and protozoa to the final point where it is resistant to further breakdown. Humus usually constitutes only five percent of the soil or less by volume but it is an essential source of nutrients and adds important textural qualities to soil critical to soil health and plant growth. Humus also hold bits of un-decomposed organic matter which feed arthropods and worms that further improve the soil. Humus has high cation exchange capacity that on a dry weight basis is many times greater than clay colloids and acts as a buffer, like clay, against changes in pH.

Humic acids and fulvic acids are important constituents of humus that begin with undecomposed organic matter. After death, these plant residues begin to decay, resulting finally in the formation of humus. With decomposition, there is a reduction of water soluble constituents including cellulose and hemicellulose; as the residues are deposited and break down, humin, lignin and lignin complexes accumulate within the soil; as microorganisms live and feed on decaying plant matter, an increase in these proteins occurs.

Lignin is resistant to breakdown and accumulates within the soil; it also chemically reacts with amino acids which add to its resistance to decomposition, including enzymatic decomposition by microbes. Fats and waxes from plant matter have some resistance to decomposition and persist in soils for a while. Clay soils often have higher organic contents that persist longer than soils without clay as the organic molecules adhere to the clay they are stabilized. Proteins normally decompose readily, but when bound to clay particles they become more resistant to decomposition. Clay particles also absorb the enzymes that would normally break down proteins. The addition of organic matter to clay soils can render that organic matter and any added nutrients inaccessible to plants and microbes for many years, since they can bind so strongly to the clay. High soil tannin (polyphenol) content from plants can cause nitrogen to be sequestered in proteins or cause nitrogen immobilization, also making nitrogen unavailable to plants.^[91][92]

Humus formation is a process dependent on the amount of plant material added each year and the type of base soil; both are affected by climate and the type of organisms present. Soils with humus can vary in nitrogen content but have 3 to 6 percent nitrogen typically; raw organic matter, as a reserve of nitrogen and phosphorus, is a vital component affecting soil fertility.^[90] Humus also absorbs water, acting as a moisture reserve that plants can utilize; it also expands and shrinks between dry and wet states, increasing soil porosity spaces. Humus is less stable than the soil's mineral constituents, as it is reduced by microbial decomposition, and over time its concentration diminshes without the addition of new organic matter. However, humus may persist over centuries if not millennia.

Climate and organics[link]

The production and accumulation or degradation of organic matter and humus is greatly dependent on climate. Temperature and soil moisture are the major factors in the formation or degradation of organic matter, they along with topography, determine the formation of organic soils. Organic matter tends to accumulate under wet or cold conditions where decomposer activity is impeded by low temperature^[93] or excess moisture which results in anaerobic conditions.^[94]

Soil horizons[link]

Horizontal layers of the soil, whose physical features, composition and age are distinct from the ones above and beneath, are referred to as soil horizons. The naming of horizons is based on the type of material of which they are composed; these materials reflect the duration of specific processes of soil formation. They are labeled using a short hand notation of letters and numbers^[95] and are described and classified by their color, size, texture, structure, consistency, root quantity, pH, voids, boundary characteristics, and presence of nodules or concretions.^[96] Few soil profiles have all the major horizons; soils may have one or many horizons.

The exposure of parent material to favorable conditions produces mineral soils that are marginally suitable for plant growth. Plant growth often results in the accumulation of organic residues. The accumulated organic layer called the O horizon produces a more active soil due to the effect of the organisms that live within it. Biological organisms colonize and break down organic materials, making available nutrients upon which other plants and animals can live. After sufficient time, humus moves downward and is deposited in a distinctive organic surface layer called the A horizon.

Classification[link]

Soil is classified into categories in order to understand relationships between different soils and to determine the suitability of a soil for a particular use. One of the first classification systems was developed by the Russian scientist Dokuchaev around 1880. It was modified a number of times by American and European researchers, and developed into the system commonly used until the 1960s. It was based on the idea that soils have a particular morphology based on the materials and factors that form them. In the 1960s, a different classification system began to emerge, that focused on soil morphology instead of parental materials and soil-forming factors. Since then it has undergone further modifications. The World Reference Base for Soil Resources (WRB)^[97] aims to establish an international reference base for soil classification.

USDA soil taxonomy[link]

A taxonomy is an arrangement in a systematic manner. Soil taxonomy has six categories. They are, from most general to specific: order, suborder, great group, subgroup, family and series. The soil properties that can be measured quantitatively are used to classify soils. A partial list is: depth, moisture, temperature, texture, structure, cation exchange capacity, base saturation, clay mineralogy, organic matter content and salt content.

In the United States, soil orders are the top hierarchical level of soil classification in the USDA soil taxonomy. the names of the orders end with the suffix -sol. There are 12 soil orders in Soil Taxonomy:^[98] The criteria for the order divisions include properties that reflect major differences in the genesis of soils.

• Alfisol - soils with aluminium and iron. They have horizons of clay accumulation, and form where there is enough moisture and warmth for at least three months of plant growth. They cover 10.1% of the soils.
• Andisols - volcanic ash soils, are young and very fertile. They cover 1% of the world's ice free surface.
• Aridisol - dry soils forming under desert conditions have fewer than 90 consecutive days of moisture during the growing season. They include nearly 12% of soils on Earth. Soil formation is slow, and accumulated organic matter is scarce. They may have subsurface zones of caliche or duripan. Many aridisols have well-developed Bt horizons showing clay movement from past periods of greater moisture.
• Entisol - recently formed soils that lack well-developed horizons. Commonly found on unconsolidated river and beach sediments of sand and clay or volcanic ash, some have an A horizon on top of bedrock. They are 18% of soils worldwide.
• Gelisols - permafrost soils with permafrost within two meters of the surface or gelic materials and permafrost within one meter. They cover 9.1% of the soils worldwide.
• Histosol - organic soils formerly called bog soils are 1.2% of soils worldwide.
• Inceptisol - young soils. They have subsurface horizon formation but show little eluviation and illuviation. They cover 15% of soils worldwide.
• Mollisol - soft, deep, dark fertile soil formed in grasslands and some hardwood forests with very thick A horizons. They are 7% of soils worldwide.
• Oxisol - are the most weathered, are rich in iron and aluminum oxides (sesquioxides) and kayolin but low in silica. They have only trace nutrients due to tropical rainfall and high temperatures. They are 7.5% of soils worldwide.
• Spodosol - acid soils with organic colloid layer complexed with iron and aluminum leached from an layer above. They are typical soils of coniferous and deciduous forests in cooler climates. They constitute 4% of soil worldwide.
• Ultisol - acid soils in humid climates, tropical to subtropical temperatures, that are heavily leached of Ca, Mg, and K nutrients. They are not quite Oxisols. They are 8.1% of the soil worldwide.
• Vertisol - inverted soils. They are clay rich and tend to swell when wet and shrink upon drying, often forming deep cracks that surface layers can fall into. They support neither farming nor construction due to their high expansion rate. They constitute 2.4% of soils worldwide.

The percentages listed above^[99] are for land area free of ice. "Soils of Mountains", which constitute the balance (11.6%), have a mixture of those listed above, or are classified as "Rugged Mountains" that have no soil.

The soil orders in sequence of increasing degree of development are Entisols, Inceptisols, Aridisols, Mollisols, Alfisols, Spodosols, Ultisols, and Oxisols. Histosols and Vertisols may appear in any of the above at any time during their development.

The soil suborders within an order are differentiated on the basis of soil properties and horizons that depend on soil moisture and temperature. Forty-seven suborders are recognized in the United States.

The soil great group category is a subdivision of a suborder. They distinguish one soil from another by the kind and sequence of soil horizons. About 185 great groups are recognized in the United States and are established on the basis of differentiating soil horizons and soil features. Horizons marked by clay, iron, humus and hard pans and soil features that are self-mixing such as clay, temperature, and marked quantities of various salts are used.

The great group categories are divided into three kinds of soil subgroups: typic, intergrade and extragrade. A typic subgroup represents the basic or "typical" concept of the great group to which the described subgroup belongs. An intergrade subgroup describes the properties that suggest how it grades (is similar to) toward soils of other soil great groups, suborders or orders. These properties are not developed or expressed well enough to include the described soil within the great group toward which they grade but suggest similarities. Extragrade features describes aberrant properties that prevent that soil from being included in another soil classification. There are about 1,000 subgroups in the United States.

A soil family category is a group of soils within a subgroup and describes the physical and chemical properties that affect the response of soil to agricultural management and engineering application. The principal characteristics used to differentiate soil families include texture, mineralogy, pH, permeability, structure, consistency, area's precipitation pattern, and soil temperature. For some soils the criteria also specify the percentage of silt, sand and coarse fragments such as gravel, cobbles and rocks. About 4,500 soil families are recognized in the United States.

A family may contain several soil series that describes the physical location by way of a name of a prominent physical feature such as a river, town, etc. near where the soil sample was taken. An example would be Merrimac for the Merrimac River in New Hampshire, USA. More than 14,000 soil series are recognized in the United States. This allows very specific descriptions to be made about soils.

Uses[link]

Soil is used in agriculture, where it serves as the anchor and primary nutrient base for plants; however, as demonstrated by hydroponics, it is not essential to plant growth if the soil-contained nutrients could be dissolved in a solution. The types of soil and available moisture determine the species of plants that can be cultivated.

Soil material is a critical component in the mining and construction industries. Soil serves as a foundation for most construction projects. The movement of massive volumes of soil can be involved in surface mining, road building and dam construction. Earth sheltering is the architectural practice of using soil for external thermal mass against building walls.

Soil resources are critical to the environment, as well as to food and fiber production. Soil provides minerals and water to plants. Soil absorbs rainwater and releases it later, thus preventing floods and drought. Soil cleans the water as it percolates through it. Soil is the habitat for many organisms: the major part of known and unknown biodiversity is in the soil, in the form of invertebrates (earthworms, woodlice, millipedes, centipedes, snails, slugs, mites, springtails, enchytraeids, nematodes, protists), bacteria, archaea, fungi and algae; and most organisms living above ground have part of them (plants) or spend part of their life cycle (insects) belowground. Above-ground and below-ground biodiversities are tightly interconnected,^[100][101] making soil protection of paramount importance for any restoration or conservation plan.

The biological component of soil is an extremely important carbon sink since about 57% of the biotic content is carbon. Even on desert crusts, cyanobacteria lichens and mosses capture and sequester a significant amount of carbon by photosynthesis. Poor farming and grazing methods have degraded soils and released much of this sequestered carbon to the atmosphere. Restoring the world's soils could offset some of the huge increase in greenhouse gases causing global warming while improving crop yields and reducing water needs.^{[102][103][104]}

Waste management often has a soil component. Septic drain fields treat septic tank effluent using aerobic soil processes. Landfills use soil for daily cover. Land application of wastewater relies on soil biology to aerobically treat BOD.

Organic soils, especially peat, serve as a significant fuel resource; but wide areas of peat production, such as sphagnum bogs, are now protected because of patrimonial interest.

Both animals and humans in many cultures occasionally consume soil. It has been shown that some monkeys consume soil, together with their preferred food (tree foliage and fruits), in order to alleviate tannin toxicity.^[105][1]

Soils filter and purify water and affect its chemistry. Rain water and pooled water from ponds, lakes and rivers percolate through the soil horizons and the upper rock strata, thus becoming groundwater. Pests (viruses) and pollutants, such as persistent organic pollutants (chlorinated pesticides, polychlorinated biphenyls), oils (hydrocarbons), heavy metals (lead, zinc, cadmium), and excess nutrients (nitrates, sulfates, phosphates) are filtered out by the soil.^[106] Soil organisms metabolize them or immobilize them in their biomass and necromass,^[107] thereby incorporating them into stable humus.^[108] The physical integrity of soil is also a prerequisite for avoiding landslides in rugged landscapes.^[109]

Degradation[link]

Here, land degradation^[110] refers to human-induced or natural process which impairs the capacity of land to function. Soils are the critical component in land degradation when it involves acidification, contamination, desertification, erosion or salination.