The Soil
Soil is the mixture of
minerals, organic matter, gases, liquids, and the myriad of organisms that together support plant life. It is a natural body that exists as part of the pedosphere
and which performs four important functions: it is a medium for plant
growth; it is a means of water storage, supply and purification; it is a
modifier of the atmosphere;
and it is a habitat for organisms that take part in decomposition of
organic matter and the creation of a habitat for new organisms.
Soil is considered to be the "skin of the earth" with interfaces between the
lithosphere, hydrosphere, atmosphere, and biosphere. Soil consists of a solid phase (minerals and organic matter) as well as a porous phase that holds gases and water. Accordingly, soils are often treated as a three-state system.
Soil is the end product of the influence of the
climate, relief
(elevation, orientation, and slope of terrain), biotic activities
(organisms), and parent materials (original minerals) interacting over
time. Soil continually undergoes development by way of numerous physical, chemical and biological processes, which include weathering with associated erosion.
Most soils have a density between 1 and 2 g/cm
3.
Little of the soil of planet
Earth is older than the Pleistocene and none is older than the Cenozoic, although fossilized soils are preserved from as far back as the Archean.
Soil science has two main branches of study:
Edaphology
and Pedology. Pedology is focused on the formation, description
(morphology), and classification of soils in their natural environment,
whereas edaphology is concerned with the influence of soils on
organisms. In engineering terms, soil is referred to as regolith, or
loose rock material that lies above the 'solid geology'.Soil is commonly referred to as "earth" or "dirt"; technically, the term "dirt" should be restricted to displaced soil.
As soil resources serve as a basis for food security, the
international community advocates for its sustainable and responsible
use through different types of
Soil Governance.
Overview
Soil is a major component of the Earth's ecosystem. From ozone
depletion and global warming to rain forest destruction and water
pollution, the world's ecosystems are impacted in far-reaching ways by
the processes carried out in the soil. Soil is the largest surficial
global carbon reservoir on Earth, and it is potentially one of the most
reactive to human disturbance and climate change. As the planet warms,
soils will add carbon dioxide to the atmosphere due to its increased
biological activity at higher temperatures. Thus, soil carbon losses
likely have a large positive feedback response to global warming.
Soil acts as an engineering medium, a habitat for soil organisms, a
recycling system for nutrients and organic wastes, a regulator of water
quality, a modifier of atmospheric composition, and a medium for plant
growth. Since soil has a tremendous range of available niches and
habitats, it contains most of the earth's genetic diversity. A gram of
soil can contain billions of organisms, belonging to thousands of species.
[citation needed] Soil has a mean prokaryotic density of roughly 10
13 organisms per cubic meter, whereas the ocean has a mean prokaryotic density of roughly 10
8 organisms per cubic meter. The carbon content stored in soil is eventually returned to the atmosphere through the process of
respiration, which is carried out by heterotrophic organisms that feed upon the carbonaceous
material in the soil. Since plant roots need oxygen, ventilation is an
important characteristic of soil. This ventilation can be accomplished
via networks of soil pores, which also absorb and hold rainwater making
it readily available for plant uptake. Since plants require a nearly
continuous supply of water, but most regions receive sporadic rainfall,
the water-holding capacity of soils is vital for plant survival.
Soils can effectively remove impurities, kill disease agents, and
degrade contaminants. Typically, soils maintain a net absorption of
oxygen and methane, and undergo a net release of carbon dioxide and nitrous oxide.
Soils offer plants physical support, air, water, temperature
moderation, nutrients, and protection from toxins. Soils provide readily
available nutrients to plants and animals by converting dead organic
matter into various nutrient forms.
Soils supply plants with mineral nutrients held in place by the clay
and humus content of the soil. For optimum plant growth, the generalized
content of soil components by volume should be roughly 50% solids (45%
mineral and 5% organic matter), and 50% voids of which half is occupied
by water and half by gas. The percent soil mineral and organic content
is typically treated as a constant, while the percent soil water and gas
content is considered highly variable whereby a rise in one is
simultaneously balanced by a reduction in the other.
The pore space allows for the infiltration and movement of air and
water, both of which are critical for life in soil. Compaction, a common
problem with soils, reduces this space, preventing air and water from
reaching the plant roots and soil organisms.
Given sufficient time, an undifferientated soil will evolve a
soil profile which consists of two or more layers, referred to as soil horizons,
that differ in one or more properties such as in their texture,
structure, density, porosity, consistency, temperature, color, and
reactivity. The horizons differ greatly in thickness and generally lack
sharp boundaries. Soil profile development is dependent on the processes
that form soils from their parent materials, the type of parent
material, and the factors that control soil formation. The biological
influences on soil properties are strongest near the surface, while the
geochemical influences on soil properties increase with depth. Mature
soil profiles in temperate climate regions typically include three basic master horizons: A, B and C. The solum normally includes the A and B horizons. The living component of the soil is largely confined to the solum. In the more hot, humid, climate of the tropics, a soil may have only a single horizon.
The
soil texture
is determined by the relative proportions of sand, silt, and clay in
the soil. The addition of organic matter, water, gases and time causes
the soil of a certain texture to develop into a larger soil structure called an aggregate.
At that point a soil can be said to be developed, and can be described
further in terms of color, porosity, consistency, reaction etc.
Of all the factors influencing the evolution of soil, water is the
most powerful due to its involvement in the solution, erosion,
transportation, and deposition of the materials of which a soil is
composed. The mixture of water and dissolved and suspended materials is
called the soil solution. Since soil water is never pure water, but
contains hundreds of dissolved organic and inorganic substances, it may
be more accurately called the soil solution. Water is central to the
solution, precipitation and leaching
of minerals from the soil profile. Finally, water affects the type of
vegetation that grows in a soil, which in turn affects the development
of the soil profile.
The most influential factor in stabilizing soil fertility are the soil
colloidal
particles, clay and humus, which behave as repositories of nutrients
and moisture and so act to buffer the variations of soil solution ions
and moisture. The contribution of soil colloids to soil nutrition are
out of proportion to their part of the soil. Colloids act to store
nutrients that might otherwise be leached from the soil or to release
those ions in response to changes of soil pH, and so, to make them available to plants.
The greatest influence on plant nutrient availability is soil pH,
which is a measure of the hydrogen ion (acid-forming) soil reactivity,
and is in turn a function of the soil materials, precipitation level,
and plant root behavior. Soil pH strongly affects the availability of
nutrients.
Most nutrients, with the exception of nitrogen, originate from
minerals. Some nitrogen originates from rain, but most of the nitrogen
available in soils is the result of
nitrogen fixation
by bacteria. The action of microbes on organic matter and minerals may
be to free nutrients for use, sequester them, or cause their loss from
the soil by their volatilisation to gases or their leaching from the
soil. The nutrients may be stored on soil colloids, and live or dead
organic matter, but they may not be accessible to plants due to extremes
of pH.
The organic material of the soil has a powerful effect on its
development, fertility, and available moisture. Following water and soil
colloids, organic material is next in importance to soil's formation
and fertility.
History of the study of soil
Studies concerning soil fertility
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 the 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 as dung."
Columella’s
"Husbandry," circa 60 A.D., advocated the use of lime and that clover
and alfalfa (green manure) should be turned under, and was used by 15
generations (450 years) under 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 European Dark Ages, Yahya Ibn_al-'Awwam’s
handbook, with its emphasis on irrigation, guided the people of North
Africa, Spain and the Middle East; a translation of this work was
finally carried to the southwest of the United States.
Experiments into what made plants grow first led to the idea that the
ash left behind when plant matter was burned was the essential element
but overlooked the role of nitrogen, which is not left on the ground
after combustion. In about 1635, the Flemish chemist
Jan Baptist van Helmont
thought he had proved water to be the essential element from his famous
five years' experiment with a willow tree grown with only the addition
of rainwater. His conclusion came from the fact that the increase in the
plant's weight had apparently been produced only by the addition of
water, with no reduction in the soil's weight. John Woodward
(d. 1728) experimented with various types of water ranging from clean
to muddy and found muddy water the best, and so he concluded that earthy
matter was the essential element. Others concluded it was humus in the
soil that passed some essence to the growing plant. Still others held
that the vital growth principal was something passed from dead plants or
animals to the new plants. At the start of the 18th century, Jethro Tull
demonstrated that it was beneficial to cultivate (stir) the soil, but
his opinion that the stirring made the fine parts of soil available for
plant absorption was erroneous.
As chemistry developed, it was applied to the investigation of soil fertility. The French chemist
Antoine Lavoisier
showed in about 1778 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. It was the French
agriculturalist Jean-Baptiste Boussingault
who by means of experimentation obtained evidence showing that the main
sources of carbon, hydrogen and oxygen for plants were the air and
water. Justus von Liebig in his book Organic Chemistry in its Applications to Agriculture and Physiology
(published 1840), asserted that the chemicals in plants must have come
from the soil and air and that to maintain soil fertility, the used
minerals must be replaced. Liebig nevertheless believed the nitrogen was
supplied from the air. The enrichment of soil with guano by the Incas
was rediscovered in 1802, by Alexander von Humboldt. This led to its mining and that of Chilean nitrate and to its application to soil in the United States and Europe after 1840.
The work of Liebig was a revolution for agriculture, and so other investigators started experimentation based on it. In England
John Bennet Lawes and Joseph Henry Gilbert worked in the Rothamsted Experimental Station,
founded by the former, and discovered that plants took nitrogen from
the soil, and that salts needed to be in an available state to be
absorbed by plants. Their investigations also produced the "superphosphate", consisting in the acid treatment of phosphate rock.
This led to the invention and use of salts of potassium (K) and
nitrogen (N) as fertilizers. Ammonia generated by the production of coke
was recovered and used as fertiliser.
Finally, the chemical basis of nutrients delivered to the soil in
manure was understood and in the mid-19th century chemical fertilisers
were applied. However, the dynamic interaction of soil and its life
forms awaited discovery.
In 1856 J. T. Way discovered that ammonia contained in fertilisers
was transformed into nitrates, and twenty years later R. W. Warington
proved that this transformation was done by living organisms. In 1890
Sergei Winogradsky announced he had found the bacteria responsible for this transformation.
It was known that certain
legumes
could take up nitrogen from the air and fix it to the soil but it took
the development of bacteriology towards the end of the 19th century to
lead to an understanding of the role played in nitrogen fixation by
bacteria. The symbiosis of bacteria and leguminous roots, and the fixation of nitrogen by the bacteria, were simultaneously discovered by German agronomist Hermann Hellriegel and Dutch microbiologist Martinus Beijerinck.
Crop rotation, mechanisation, chemical and natural fertilisers led to a doubling of wheat yields in Western Europe between 1800 and 1900.
Studies concerning soil formation
The scientists who studied the soil in connection with agricultural
practices had considered it mainly as a static substrate. However, soil
is the result of evolution from more ancient geological materials. After
studies of the improvement of the soil commenced, others began to study
soil genesis and as a result also soil types and classifications.
In 1860, in Mississippi,
Eugene W. Hilgard
studied the relationship among rock material, climate, and vegetation,
and the type of soils that were developed. He realised that the soils
were dynamic, and considered soil types classification. Unfortunately
his work was not continued. At the same time Vasily Dokuchaev
(about 1870) was leading a team of soil scientists in Russia who
conducted an extensive survey of soils, finding that similar basic
rocks, climate and vegetation types lead to similar soil layering and
types, and established the concepts for soil classifications. Due to the
language barriers, the work of this team was not communicated to
Western Europe until 1914 by a publication in German by K. D. Glinka, a
member of the Russian team.
Curtis F. Marbut
was influenced by the work of the Russian team, translated Glinka's
publication into English, and as he was placed in charge of the U. S. National Cooperative Soil Survey, applied it to a national soil classification system.
Soil-forming processes
Soil formation, or
pedogenesis, is the combined effect of physical, chemical, biological and anthropogenic processes working on soil parent material. Soil is said to be formed when organic matter has accumulated and colloids
are washed downward, leaving deposits of clay, humus, iron oxide,
carbonate, and gypsum, producing a distinct layer called the B horizon.
(This is a somewhat arbitrary definition as mixtures of sand, silt, clay
and humus will support biological and agricultural activity before that
time.) These constituents are moved from one level to another by water
and animal activity. As a result, layers (horizons) form in the soil
profile. The alteration and movement of materials within a soil causes
the formation of distinctive soil horizons.
How soil formation proceeds is influenced by at least five classic
factors that are intertwined in the evolution of a soil. They are:
parent material,
climate,
topography (relief),
organisms, and
time. When reordered to climate, relief, organisms, parent material, and time, they form the acronym CROPT.
An example of the development of a soil would begin with the
weathering of lava flow bedrock, which would produce the purely
mineral-based parent material from which the soil texture forms. Soil
development would proceed most rapidly from bare rock of recent flows in
a warm climate, under heavy and frequent rainfall. Under such
conditions, 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 the rocks and guano.
Crevasses and pockets, local topography of the rocks, would hold fine
materials and harbour plant roots. The developing plant roots are
associated with mycorrhizal fungi that assist in breaking up the porous lava, and by these means organic matter and a finer mineral soil accumulate with time.
Parent material
The mineral material from which a soil forms is called
parent material.
Rock, whether its origin is igneous, sedimentary, or metamorphic, is
the source of all soil mineral materials and the origin of all plant
nutrients with the exceptions of nitrogen, hydrogen and carbon. As the
parent material is chemically and physically weathered, transported, deposited and precipitated, it is transformed into a soil.
Typical soil parent mineral materials are:
- Quartz: SiO2
- Calcite: CaCO3
- Feldspar: KAlSi3O8
- Mica (biotite): K(Mg,Fe)3AlSi3O10(OH)2
Classification of parent material
- Parent materials are classified according to how they came to be deposited. Residual materials are mineral materials that have weathered in place from primary bedrock. Transported materials are those that have been deposited by water, wind, ice or gravity. Cumulose material is organic matter that has grown and accumulates in place.
- Residual soils are soils that develop from their underlying
parent rocks and have the same general chemistry as those rocks. The
soils found on mesas, plateaux, and plains are residual soils. In the
United States as little as three percent of the soils are residual.
- Most soils derive from transported materials that have been moved many miles by wind, water, ice and gravity.
-
- Aeolian processes (movement by wind) are capable of moving silt and fine sand many hundreds of miles, forming loess soils (60–90 percent silt), common in the Midwest of North America and in Central Asia. Clay is seldom moved by wind as it forms stable aggregates.
-
- Water-transported materials are classed as either alluvial, lacustrine, or marine. Alluvial materials are those moved and deposited by flowing water. Sedimentary deposits settled in lakes are called lacustrine. Lake Bonneville
and many soils around the Great Lakes of the United States are
examples. Marine deposits, such as soils along the Atlantic and Gulf
Coasts and in the Imperial Valley of California of the United States,
are the beds of 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 is not moved but originates from deposited organic material. This includes peat and muck soils
and results from preservation of plant residues by the low oxygen
content of a high water table. While peat may form sterile soils, muck
soils may be very fertile.
Weathering of parent material
- The weathering of parent material takes the form of physical weathering (disintegrating), chemical weathering (decomposition) and chemical transformation.
Generally, minerals that are formed under the high temperatures and
pressures at great depths within the earth's mantle are less resistant
to weathering, while minerals formed at low temperature and pressure
environment of the surface are more resistant to weathering. Weathering
is usually confined to the top few meters of geologic material, because
physical, chemical, and biological stresses generally decrease with
depth. Physical disintegration begins as rocks that have solidified deep
in the earth are exposed to lower pressure near the surface and swell
and become unstable. Chemical decomposition is a function of mineral
solubility, the rate of which doubles with each 10°C rise in
temperature, but is strongly dependent on water to effect chemical
changes. Rocks that will decompose in a few years in tropical climates
will remain unaltered for millennia in deserts. Structural changes are the result of hydration, oxidation, and reduction.
-
- Physical disintegration is the first stage in the
transformation of parent material into soil. Temperature fluctuations
cause expansion and contraction of the rock, splitting it along lines of
weakness. Water may then enter the cracks and freeze and cause the
physical splitting of material along a path toward the center of the
rock, while temperature gradients within the rock can cause exfoliation
of "shells". Cycles of wetting and drying cause soil particles to be
abraded to a finer size, as does the physical rubbing of material as it
is moved by wind, water, and gravity. Water can deposit within rocks
minerals that expand upon drying, thereby stressing the rock. Finally,
organisms reduce parent material in size through the action of plant
roots or digging on the part of animals
-
- Chemical decomposition and structural changes result
when minerals are made soluble by water or are changed in structure. The
first three of the following list are solubility changes and the last
three are structural changes.
-
- The solution of salts in water results from the action of bipolar water on ionic salt compounds producing a solution of ions and water.
- Hydrolysis
is the transformation of minerals into polar molecules by the splitting
of the intervening water. This results in soluble acid-base pairs. For
example, the hydrolysis of orthoclase-feldspar transforms it to acid
silicate clay and basic potassium hydroxide, both of which are more
soluble.
- In carbonation,
the reaction of carbon dioxide in solution with water forms carbonic
acid. Carbonic acid will transform calcite into more soluble calcium
bicarbonate.
- Hydration is the inclusion of water in a mineral structure, causing it to swell and leaving it more stressed and easily decomposed.
- Oxidation
of a mineral compound is the inclusion of oxygen in a mineral, causing
it to increase its oxidation number and swell due to the relatively
large size of oxygen, leaving it stressed and more easily attacked by
water (hydrolysis) or carbonic acid (carbonation).
- Reduction
the opposite of oxidation, means the removal of oxygen, hence oxidation
number of some part of the mineral is reduced, which occurs when oxygen
is scarce. The reduction of minerals leaves them electrically unstable,
more soluble and internally stressed and easily decomposed.
- Of the above, hydrolysis and carbonation are the most effective.
- 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, chelation from organic compounds, hydration
(the solution of minerals in water with resulting cation and anion
pairs) and physical processes that include freezing and thawing.
The mineralogical and chemical composition of the primary bedrock
material, its physical features, including grain size and degree of
consolidation, and the rate and type of weathering transform the parent
material into a different mineral. The texture, pH and mineral
constituents of saprolite are inherited from its parent material.
Climate
The principal climatic variables influencing soil formation are
effective precipitation (i.e., precipitation minus evapotranspiration)
and temperature, both of which affect the rates of chemical, physical,
and biological processes. The temperature and moisture both influence
the organic matter content of soil through their effects on the balance
between plant growth and microbial decomposition. Climate is the
dominant factor in soil formation, and soils show the distinctive
characteristics of the
climate zones in which they form. For every 10 °C rise in temperature, the rates of biochemical reactions more than double.
Mineral precipitation and temperature are the primary climatic
influences on soil formation. If warm temperatures and abundant water
are present in the profile at the same time, the processes of
weathering, leaching, and plant growth will be maximized. Humid climates
favor the growth of trees. In contrast, grasses are the dominant native
vegetation in subhumid and semiarid regions, while shrubs and brush of
various kinds dominate in arid areas.
Water is essential for all the major chemical weathering reactions.
To be effective in soil formation, water must penetrate the regolith.
The seasonal rainfall distribution, evaporative demand, site topography,
and soil permeability interact to determine how effectively
precipitation can influence soil formation. The greater the depth of
water penetration, the greater the depth of weathering of the soil and
its development. Surplus water percolating through the soil profile
transports soluble and suspended materials from the upper to the lower
layers. It may also carry away soluble materials in the drainage waters.
Thus, percolating water stimulates weathering reactions and helps
differentiate soil horizons. Likewise, a deficiency of water is a major
factor in determining the characteristics of soils of dry regions.
Soluble salts are not leached from these soils, and in some cases they
build up to levels that curtail plant growth. Soil profiles in arid and
semi-arid regions are also apt to accumulate carbonates and certain
types of expansive clays.
The direct influences of climate include:
- A shallow accumulation of lime in low rainfall areas as caliche
- Formation of acid soils in humid areas
- Erosion of soils on steep hillsides
- Deposition of eroded materials downstream
- 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 in wet and 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 which breaks up rocks and other consolidated materials.
Climate also indirectly influences soil formation through the effects
of vegetation cover and biological activity, which modify the rates of
chemical reactions in the soil.
Topography
The topography, or
relief,
is characterized by the inclination (slope), elevation, and orientation
of the terrain. Topography determines the rate of precipitation or
runoff and rate of formation or erosion of the surface soil profile. The
topographical setting may either hasten or retard the work of climatic
forces.
Steep slopes encourage rapid soil loss by erosion and allow less
rainfall to enter the soil before running off and hence, little mineral
deposition in lower profiles. In semiarid regions, the lower effective
rainfall on steeper slopes also results in less complete vegetative
cover, so there is less plant contribution to soil formation. For all of
these reasons, steep slopes prevent the formation of soil from getting
very far ahead of soil destruction. Therefore, soils on steep terrain
tend to have rather shallow, poorly developed profiles in comparison to
soils on nearby, more level sites.
In swales and depressions where runoff water tends to concentrate,
the regolith is usually more deeply weathered and soil profile
development is more advanced. However, in the lowest landscape
positions, water may saturate the regolith to such a degree that
drainage and aeration are restricted. Here, the weathering of some
minerals and the decomposition of organic matter are retarded, while the
loss of iron and manganese is accelerated. In such low-lying
topography, special profile features characteristic of wetland soils may
develop. 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
Soil is the most abundant
ecosystem on Earth, but the vast majority of organisms in soil are microbes, a great many of which have not been described.
There may be a population limit of around one billion cells per gram of
soil, but estimates of the number of species vary widely. Estimates range from over a million species per gram of soil to 50,000 per gram of soil. The total number of organisms and species can vary widely according to soil type, location, and depth.
Plants, animals, fungi,
bacteria and humans affect soil formation (see soil biomantle and stonelayer). Animals, soil mesofauna 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 metres 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 impact soil formation by removing vegetation cover with
erosion as the result. Their tillage also mixes the different soil
layers, restarting the soil formation process as less weathered material
is mixed with the more developed upper layers.
Earthworms, ants and termites mix the soil as they burrow,
significantly affecting soil formation. Earthworms ingest soil particles
and organic residues, enhancing the availability of plant nutrients in
the material that passes through their bodies. They aerate and stir the
soil and increase the stability of soil aggregates, thereby assuring
ready infiltration of water. As they build mounds, some organisms might
transport soil materials from one horizon to another.
In general, the mixing activities of animals, sometimes called
pedoturbation, tends to undo or counteract the tendency of other
soil-forming processes that create distinct horizons. Termites and ants
may also retard soil profile development by denuding large areas of soil
around their nests, leading to increased loss of soil by erosion. Large
animals such as gophers, moles, and prairie dogs bore into the lower
soil horizons, bringing materials to the surface. Their tunnels are
often open to the surface, encouraging the movement of water and air
into the subsurface layers. In localized areas, they enhance mixing of
the lower and upper horizons by creating, and later refilling,
underground tunnels. Old animal burrows in the lower horizons often
become filled with soil material from the overlying A horizon, creating
profile features known as crotovinas.
Vegetation impacts soils in numerous ways. It can prevent erosion
caused by excessive rain that might result from surface runoff. Plants
shade soils, keeping them cooler and slow 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 the soil structure. The type
and amount of vegetation depends on climate, 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.
Human activities widely influence soil formation. For example, it is
believed that Native Americans regularly set fires to maintain several
large areas of prairie grasslands in Indiana and Michigan. In more
recent times, human destruction of natural vegetation and subsequent
tillage of the soil for crop production has abruptly modified soil
formation. Likewise, irrigating an arid region of soil drastically
influences the soil-forming factors, as does adding fertilizer and lime
to soils of low fertility.
Time
Time is a factor in the interactions of all the above. While a
mixture of sand, silt and clay constitute the texture of a soil and the
aggregation of those components produces peds, the development of a soil
with a distinct B horizon marks the development of a soil. With time,
soils will evolve features that depend on the interplay of the prior
listed soil-forming factors. It takes decades to several thousand years
for a soil to develop a profile. That time period depends strongly on
climate, parent material, relief, and biotic activity.
For example, recently deposited material from a flood exhibits no soil
development as 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
time the soil will develop a profile that depends on the intensities of
biota and climate. While a 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.
[citation needed]
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 or 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
depends on climate, topography and biological activity.
Physical properties of soils
The physical properties of soils, in order of decreasing importance, are
texture,
structure,
density,
porosity,
consistency,
temperature,
colour and
resistivity.
Soil texture is determined by the relative proportion of the three
kinds of soil particles, called soil separates: sand, silt, and clay. At
the next larger scale, soil structures called
peds
are created from the soil separates when iron oxides, carbonates, clay,
silica and humus, coat particles and cause them to adhere into larger,
relatively stable secondary structures. Soil density, particularly bulk
density, is a measure of soil compaction. Soil porosity consists of the
void part of the soil volume and is occupied by gases or water. Soil
consistency is the ability of soil to stick together. Soil temperature
and colour are self-defining. Resistivity refers to the resistance to
conduction of electric currents and affects the rate of corrosion of
metal and concrete structures. These properties may vary through the
depth of a soil profile. Most of these properties determine the aeration
of the soil and the ability of water to infiltrate and to be held
within the soil.
Generalized Influence of Soil Texture Separates on Some Properties/Behavior of Soils
| Property/behavior |
Sand |
Silt |
Clay |
| Water-holding capacity |
Low |
Medium to high |
High |
| Aeration |
Good |
Medium |
Poor |
| Drainage rate |
High |
Slow to medium |
Very slow |
| Soil organic matter level |
Low |
Medium to high |
High to medium |
| Decomposition of organic matter |
Rapid |
Medium |
Slow |
| Warm-up in spring |
Rapid |
Moderate |
Slow |
| Compactability |
Low |
Medium |
High |
| Susceptibility to wind erosion |
Moderate (High if fine sand) |
High |
Low |
| Susceptibility to water erosion |
Low (unless fine sand) |
High |
Low if aggregated, otherwise high |
| Shrink/Swell Potential |
Very Low |
Low |
Moderate to very high |
| Sealing of ponds, dams, and landfills |
Poor |
Poor |
Good |
| Suitability for tillage after rain |
Good |
Medium |
Poor |
| Pollutant leaching potential |
High |
Medium |
Low (unless cracked) |
| Ability to store plant nutrients |
Poor |
Medium to High |
High |
| Resistance to pH change |
Low |
Medium |
High |
Texture
Main article:
Soil texture
The mineral components of soil are sand, silt and clay, and their
relative proportions determine a soil's texture. Properties that are
influenced by soil texture, include porosity, permeability,
infiltration, shrink-swell rate, water-holding capacity, and
susceptibility to erosion. In the illustrated
USDA textural classification triangle, the only soil in which neither sand, silt nor clay predominates 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 behaviour, in particular its retention
capacity for nutrients and water.
Sand and silt are the products of physical and chemical weathering of
the parent rock; clay, on the other hand, is a product of the
precipitation of the dissolved parent rock as a secondary mineral. It is
the large surface area to volume ratio (
specific surface area) of soil particles and the unbalanced ionic charges within those that determine their role in the cation exchange capacity
of soil, and hence its fertility. Sand is least active, followed by
silt; clay is the most active. Sand's greatest benefit to soil is that
it resists compaction and increases a soil's porosity. Silt is
mineralogically like sand but with its higher specific surface area it
is more chemically active than sand. But it is the clay content of soil,
with its very high specific surface area and generally large number of
negative charges, that gives a soil its high retention capacity for
water and nutrients. Clay soils also resist wind and water erosion
better than silty and sandy soils, as the particles bond tightly 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 to 0.05 mm (0.0787 to 0.0020 in) in diameter. Silt ranges in
size from 0.05 to 0.002 mm (0.002 to 0.00008 in). Clay cannot be
resolved by optical microscopes as its particles are 0.002 mm (7.9
×10
−5 in) or less in diameter. 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 (0.079 in) 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 gravelly 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:
- Mineral fraction is 0% clay and organic matter is 20% or more
- Mineral fraction is 0% to 50% clay and organic matter is between 20% and 30%
- Mineral fraction is 50% or more clay and organic matter 30% or more.
Structure
The clumping of the soil textural components of sand, silt and clay causes
aggregates to form and the further association of those aggregates into larger units creates
soil structures called peds.
The adhesion of the soil textural components by organic substances,
iron oxides, carbonates, clays, and silica, and the breakage of those
aggregates from expansion-contraction, freezing-thawing, and
wetting-drying cycles, shape soil into distinct geometric forms. The
peds evolve into units which may have various shapes, sizes and degrees of development.
A soil clod, however, is not a ped but rather a mass of soil that
results from mechanical disturbance of the soil. Soil structure affects
aeration, water movement, conduction of heat, plant root growth and
resistance to erosion. Water, in turn, has its 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 the
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 that does not change with agricultural activities,
soil structure can be improved or destroyed by the choice and timing of
farming practices.
Soil Structural Classes:
- 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 textural constituents, 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, due to its
differential drying rate with respect to the surface, will induce
horizontal cracks, reducing columns to blocky peds. Roots, rodents,
worms, and freezing-thawing cycles further break the peds into a
spherical shape.
At a smaller scale, plant roots extend into voids and remove water
causing the open spaces to increase, and decrease physical aggregation
size. At the same time roots, fungal
hyphae and earthworms create microscopic tunnels that break up peds.
At an even smaller scale, soil aggregation continues as bacteria and
fungi exude sticky polysaccharides which 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 which give the 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
(form clumps). 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 the
particles to push apart, and by doing so to prevent the flocculation of
clay particles into larger, open assemblages. As a result, the clay
disperses and settles into voids between peds, causing those to close.
In this way the soil aggregation is destroyed and the soil is made
impenetrable to air and water. Such sodic soil tends to form columnar structures near the surface.
Density
Density is the weight per unit volume of an object. Particle density
is equal to the mass of solid particles divided by the volume of solid
particles - it is the density of only the mineral particles that make up
a soil; i.e., it excludes pore space and organic material. Soil
particle density is typically 2.60 to 2.75 grams per cm
3 and
is usually unchanging for a given soil. Soil particle density is lower
for soils with high organic matter content, and is higher for soils with
high Fe-oxides content. Soil bulk density is equal to the dry mass of
the soil divided by the volume of the soil; i.e., it includes air space
and organic materials of the soil volume. A high bulk density is
indicative of either soil compaction or high sand content. The bulk
density of cultivated loam is about 1.1 to 1.4 g/cm
3 (for comparison water is 1.0 g/cm
3).
Soil bulk density is highly variable for a given soil. A lower bulk
density by itself does not indicate suitability for plant growth due to
the influence of soil texture and structure. Soil bulk density is
inherently always less than the soil particle density.
Representative bulk densities of soils. The percentage pore space was calculated using 2.7 g/cm3 for particle density except for the peat soil, which is estimated.
| Soil treatment and identification |
Bulk density g/cm3 |
Pore space % |
| Tilled surface soil of a cotton field |
1.3 |
51 |
| Trafficked inter-rows where wheels passed surface |
1.67 |
37 |
| Traffic pan at 25 cm deep |
1.7 |
36 |
| Undisturbed soil below traffic pan, clay loam |
1.5 |
43 |
| Rocky silt loam soil under aspen forest |
1.62 |
40 |
| Loamy sand surface soil |
1.5 |
43 |
| Decomposed peat |
0.55 |
65 |
Porosity
Pore space is that part of the bulk volume of soil that is not
occupied by either mineral or organic matter but is open space occupied
by either gases or water. Ideally, the total pore space should be 50% of
the soil volume. The gas 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. This
property of soils effectively compartmentalizes the soil pore space such
that many organisms are not in direct competition with one another,
which may explain not only the large number of species present, but the
fact that functionally redundant organisms (organisms with the same
ecological niche) can co-exist within the same soil.
There are four categories of pores:
- Very fine pores: < 2 µm
- Fine pores: 2-20 µm
- Medium pores: 20-200 µm
- Coarse pores: 200 µm-0.2 mm
In comparison, root hairs are 8 to 12 µm in diameter. When pore space
is less than 30 µm, the forces of attraction that hold water in place
are greater than the gravitational force 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 gas and water movement is superior
to smaller pore space soil that has a greater percentage pore space.
Soil texture determines the pore space at the smallest scale, but at a
larger scale, soil structure has a strong influence on soil aeration,
water infiltration and drainage.
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
. Clay soils have smaller pores, but more total pore space than sand.
Consistency
Consistency is the ability of soil to stick to itself or to other
objects (cohesion and adhesion respectively) and its ability to resist
deformation and rupture. It is of approximate use in predicting
cultivation problems and the engineering of foundations. Consistency is
measured at three moisture conditions: air-dry, moist, and wet; In those
conditions the consistency quality depend upon the clay content. In the
wet state, the two qualities of stickiness and plasticity are assessed.
A soil's resistance to fragmentation and crumbling is assessed in the
dry state by rubbing the sample. Its resistance to shearing forces is
assessed in the moist state by thumb and finger pressure. Finally, a
soil's plasticity is measured in the wet state by moulding with the
hand. Finally, the cemented consistency depends on cementation by
substances other than clay, such as calcium carbonate, silica, oxides
and salts; moisture content has little effect on its assessment. The
measures of consistency border on subjective as they employ the "feel"
of the soil in those states.
The terms used to describe the soil consistency in three moisture
states and a last consistency not affected by the amount of moisture are
as follows:
- Consistency of Dry Soil: loose, soft, slightly hard, hard, very hard, extremely hard
- Consistency of Moist Soil: loose, very friable, friable, firm, very firm, extremely firm
- Consistency of Wet Soil: nonsticky, slightly sticky, sticky, very sticky; nonplastic, slightly plastic, plastic, very plastic
- Consistency of Cemented Soil: weakly cemented, strongly cemented, indurated (requires hammer blows to break up)
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
Soil temperature depends on the ratio of the energy absorbed to that
lost. Soil has a temperature range between -20 to 60 °C. Soil
temperature regulates seed germination, plant and root growth and the
availability of nutrients. Below 50 cm (20 in), soil temperature seldom
changes and can be approximated by adding 1.8 °C (2 °F) to the mean
annual air temperature. Soil temperature has important seasonal, monthly
and daily variations. Fluctuations in soil temperature are much lower
with increasing soil depth. Heavy
mulching (a type of soil cover) can slow the warming of soil, and, at the same time, reduce fluctuations in surface temperature.
Most often, agricultural activities must adapt to soil temperatures by:
- maximizing germination and growth by timing of planting
- optimizing use of anhydrous ammonia by applying to soil below 10 °C (50 °F)
- preventing heaving and thawing due to frosts from damaging shallow-rooted crops
- preventing damage to desirable soil structure by freezing of saturated soils
- improving uptake of phosphorus by plants
Otherwise soil temperatures can be raised by drying soils or the use
of clear plastic mulches. Organic mulches slow the warming of the soil.
There are various factors that affect soil temperature, such as water
content, soil color, and relief (slope, orientation, and elevation),
and soil cover (shading and insulation). The color of the ground cover
and its insulating properties have a strong influence on soil
temperature. Whiter soil tends to have a higher
albedo
than blacker soil cover, which encourages whiter soils to have cooler
soil temperatures. The specific heat of soil is the energy required to
raise the temperature of soil by 1 °C. The specific heat of soil
increases as water content increases, since the heat capacity of water
is greater than that of dry soil. The specific heat of pure water is ~ 1
calorie per gram, the specific heat of dry soil is ~ 0.2 calories per
gram and the specific heat of wet soil is ~ 0.2 to 1 calories per gram.
Also, tremendous energy (~540 cal/g) is required and dissipated to
evaporate water (known as the heat of vaporization).
As such, wet soil usually warms more slowly than dry soil - wet surface
soil is typically 3 to 6 °C colder than dry surface soil.
Soil heat flux refers to the conduction (or movement) of energy (or
heat) in response to a temperature gradient. The heat flux density is
the amount of energy that flows through soil per unit area per unit time
has both magnitude and direction.
where (including the
SI units)
is the local heat flux, W·m−2
is the material's conductivity, W·m−1·K−1,
where the del operator applied to the temperature gives the temperature gradient, K·m−1 and has both magnitude and direction.
The thermal conductivity,
,
is often treated as a constant, though this is not always true. While
the thermal conductivity of a material generally varies with
temperature, the variation is generally small over a significant range
of temperatures for some common materials. In anisotropic materials, the
thermal conductivity typically varies with orientation; in this case
is represented by a second-order
tensor. In nonuniform materials,
varies with spatial location. For soil, thermal conductivity also
depends on mineral composition, water content, and bulk density. Compact
and wet soils have a higher thermal conductivity than loose and dry
soils. For many simple applications, Fourier's law is used in its
one-dimensional, x-direction form:
| Component |
Thermal Conductivity (W.m‐1.K‐1) |
| Quartz |
8.8 |
| Clay |
2.9 |
| Organic Matter |
0.25 |
| Water |
0.57 |
| Ice |
2.4 |
| Air |
0.025 |
| Dry soil |
0.2‐0.4 |
| Wet soil |
1-3 |
Colour
Soil colour is often the first impression one has when viewing soil.
Striking colours 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 of North America 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 the organic matter content,
drainage conditions, and degree of oxidation. Soil color, while easily
discerned, has little use in predicting soil characteristics.
"Arizona Master Gardener Manual". Cooperative Extension, College of Agriculture, University of Arizona. p. Chapter 2, pp 4–8. Retrieved 27 May 2013.
It is of use in distinguishing boundaries within a soil profile,
determining 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
colours are due to various iron minerals. The development and
distribution of colour 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 colourful compounds. Iron forms secondary minerals of a yellow or red colour, organic matter decomposes into black and brown compounds, and manganese, sulfur and nitrogen can form black mineral deposits. These pigments can produce various colour patterns within a soil. Aerobic conditions produce uniform or gradual colour changes, while reducing environments (anaerobic) result in rapid colour flow with complex, mottled patterns and points of colour concentration.
Resistivity
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 resistivity and increase conductivity, thereby increasing the rate of corrosion. Soil resistivity values typically range from about 2 to 1000 Ω·m, but more extreme values are not unusual.
Soil water
Water affects soil formation, structure, stability and erosion but is
of primary concern with respect to plant growth. Water is essential to
plants for four reasons:
- It constitutes 80%-95% of the plant's protoplasm.
- It is essential for photosynthesis.
- It is the solvent in which nutrients are carried to, into and throughout the plant.
- It provides the turgidity by which the plant keeps itself in proper position.
In addition, water alters the soil profile by dissolving and
re-depositing 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, gas one-quarter of the volume,
and water one-quarter of the volume of which only half will be available
to most plants.
A flooded field will drain the
gravitational water under the influence of gravity until water's
adhesive and
cohesive forces resist further drainage at which point it has finally reached
field capacity. At that point, plants must apply
suction to draw water from a soil. When soil becomes too dry, the
available water is used up and the remaining moisture is
unavailable water
as the plant cannot produce sufficient suction to draw the water in. A
plant must produce suction that increases from zero for a flooded field
to 1/3 bar at
field dry condition. At 15 bar suction,
wilting percent, seeds will not germinate,
plants begin to wilt and then die. Water moves in soil under the
influence of gravity, osmosis and capillarity. When water enters the
soil, it displaces air from some of the pores, since air content of a
soil is inversely related to its water content.
The rate at which a soil can absorb water depends on the soil and its
other conditions. As a plant grows, its roots remove water from the
largest pores first. Soon the larger pores hold only air, and the
remaining water is found only in the intermediate- and smallest-sized
pores. The water in the smallest pores is so strongly held on particle
surfaces that plant roots cannot pull it away. Consequently, not all
soil water is available to plants. When saturated, the soil may lose
nutrients as the water drains. Water moves in a drained field under the
influence of pressure where the soil is locally saturated and by
capillarity pull. Most plant water needs are supplied from the suction
caused by of evaporation from plant leaves and 10% is supplied by
"suction" created by osmotic pressure differences between the plant
interior and the soil water. Plant roots must seek out water.
Insufficient water will damage the yield of a crop. Most of the
available water is used in transpiration to pull nutrients into the
plant.
Water retention forces
Water is retained in a soil when the
adhesive force of attraction of water's hydrogen atoms for the oxygen of soil particles and the
cohesive forces water's hydrogen feels for other water's oxygen atoms are stronger than the forces that might pull it from the soil.
When a field is flooded, the air in the soil pore 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 gases.
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 termed
suction and usually expressed in units of bars (10
5
pascal) which is just a little less than one-atmosphere pressure.
Alternatively, the terms "tension" or "moisture potential" may be used.
Moisture classification
The forces with which water is held in soils determine 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 initially able to draw in water 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
attempting to use the water. Cacti, for example, can 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 said to be
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 a plant must produce 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 at 10,000 bar suction. All water below wilting percentage is called
unavailable water.
Soil moisture content
When the soil moisture content is optimal for plant growth, the water
in the large and intermediate size pores can move about in the soil and
be easily used by plants. The amount of water remaining in a soil
drained to field capacity and the amount that is available are functions
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 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.
Wilting point, field capacity, and available water capacity of various soil textures
| Soil Texture |
Wilting Point |
Field Capacity |
Available water capacity |
| Water per foot of soil depth |
Water per foot of soil depth |
Water per foot of soil depth |
| % |
in. |
% |
in. |
% |
in. |
| Medium sand |
1.7 |
0.3 |
6.8 |
1.2 |
5.1 |
0.9 |
| Fine sand |
2.3 |
0.4 |
8.5 |
1.5 |
6.2 |
1.1 |
| Sandy loam |
3.4 |
0.6 |
11.3 |
2.0 |
7.9 |
1.4 |
| Fine sandy loam |
4.5 |
0.8 |
14.7 |
2.6 |
10.2 |
1.8 |
| Loam |
6.8 |
1.2 |
18.1 |
3.2 |
11.3 |
2.0 |
| Silt loam |
7.9 |
1.4 |
19.8 |
3.5 |
11.9 |
2.1 |
| Clay loam |
10.2 |
1.8 |
21.5 |
3.8 |
11.3 |
2.0 |
| Clay |
14.7 |
2.6 |
22.6 |
4.0 |
7.9 |
1.4 |
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
Water moves through soil due to the force of gravity, osmosis and
capillarity. At zero to one-third bar suction, water is pushed through
soil from the point of its application under the force of gravity and
the pressure gradient created by the pressure of the water; this is
called
saturated flow. At higher suction, water movement is
pulled by capillarity from wetter toward dryer soil. This is caused by
water's adhesion to soil solids, and is called
unsaturated flow.
Water infiltration and movement in soil is controlled by six factors:
- Soil texture
- Soil structure. Fine-textured soils with granular structure are most favourable to infiltration of water.
- 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.
- Depth of soil to impervious layers such as hardpans or bedrock
- The amount of water already in the soil
- Soil temperature. Warm soils take in water faster while frozen soils
may not be able to absorb depending on the type of freezing.
Water infiltration rates range from 0.25 cm (0.098 in) per hour for
high clay soils to 2.5 cm (0.98 in) per hour for sand and well
stabilised and aggregated soil structures. Water flows through the ground unevenly, called "gravity fingers", because of the surface tension between water particles.
Tree roots create paths for rainwater flow through soil by breaking
though soil including clay layers: one study showed roots increasing
infiltration of water by 153% and another study showed an increase by 27
times.
Flooding temporarily increases soil permeability in river beds, helping to
recharge aquifers.
Saturated flow
- Water applied to a soil is pushed by pressure gradients from the
point of its application where it is saturated locally to less saturated
areas. 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. In order of decreasing solubility, the leached
nutrients are:
-
- Calcium
- Magnesium, Sulfur, Potassium; depending upon soil composition
- Nitrogen; usually little, unless nitrate fertiliser was applied recently
- Phosphorus; very little as its forms in soil are of low solubility.
- 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.
Unsaturated flow
- At suctions less than one-third bar, water moves in all directions via unsaturated flow
at a rate that is dependent on the square of the diameter of the
water-filled pores. Water is pulled by capillary action due to the
adhesion force of water to the soil solids, producing a suction gradient
from wet towards 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 is slowest in fine-textured soils such as
clay.
Water uptake by plants
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 the plant's roots to its leaves. In addition, the
high concentration of salts within plant roots creates an
osmotic pressure
gradient that pushes soil water into the roots. Osmotic absorption
becomes more important during times of low water transpiration caused by
lower temperatures (for example at night) or high humidity. It is the
process that causes guttation.
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 a total of 385 miles in
length and 2,550 square feet in surface area and 14 billion hair roots
of 6,600 miles total length and 4,320 square feet total area, for a
total surface area of 6,870 square feet (83 ft squared). The total
surface area of the loam soil was estimated to be 560,000 square feet.
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 (one inch) 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 causing 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, its production
was reduced by 34%.
Consumptive use and water efficiency
Only a small fraction (0.1% to 1%) of the water used by a plant is
held within the plant. The majority is ultimately lost via
transpiration, 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.
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. Transpiration ratios for crops range from
300 to 700. For example alfalfa may have a transpiration ratio of 500
and as a result 500 kilograms of water will produce one kilogram of dry
alfalfa.
Soil atmosphere
The atmosphere of soil is radically different from the atmosphere
above. The consumption of oxygen, by microbes and plant roots and their
release of carbon dioxide, decrease oxygen and increase carbon dioxide
concentration. Atmospheric CO
2 concentration is 0.04%, but in the soil pore space it may range from 10 to 100 times that level. At extreme levels CO
2
is toxic. In addition, the soil voids are saturated with water vapour.
Adequate porosity is necessary not just to allow the penetration of
water but also to allow gases to diffuse in and out. Movement of gases
is by diffusion from high concentrations to lower. Oxygen diffuses in
and is consumed and excess levels of carbon dioxide, diffuse out with
other gases as well as water. Soil texture and structure strongly affect
soil porosity and gas diffusion. It is the total pore space (porosity)
of soil not the pore size that determines the rate of diffusion of gases
into and out of soil. A
Platy
soil structure and compacted soils (low porosity) impede gas flow, and a
deficiency of oxygen may encourage anaerobic bacteria to reduce nitrate
NO3 to the gases N2, N2O, and NO, which are 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.
Composition of soil particles
Soil particles can be classified by their chemical composition (
mineralogy) as well as their size. The particle size distribution of a soil, its texture,
determines many of the properties of that soil, but the mineralogy of
those particles can strongly modify those properties. The mineralogy of
the finest soil particles, clay, is especially important.
Gravel, sand and silt
Gravel, sand and silt are the larger soil particles, and their mineralogy is often inherited from the parent material of the soil, but may include products of weathering (such as concretions of calcium carbonate or iron oxide), or residues of plant and animal life (such as silica phytoliths). Quartz is the most common mineral in the sand or silt fraction as it is resistant to chemical weathering; other common minerals are felspars, micas and ferromagnesian minerals such as pyroxenes, amphiboles and olivines.
Mineral colloids; soil clays
Due to its high specific surface area and its unbalanced negative charges,
clay
is the most active mineral component of soil. It is a colloidal and
most often a crystalline material. In soils, clay is defined in a
physical sense as any mineral particle less than 2 µm (8
×10
−5 in)
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 as gypsum, carbonates, or quartz, are small enough
to be classified physically as clay but chemically they do not afford
the same utility as do clay minerals.
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 that is dependent on but 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. Clay minerals continue to be formed as long as the soil
exists.
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 a 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 aluminium and silicon by way
of ionic bonds that together form a single layer of clay. The spatial
arrangement of the oxygen atoms determines clay's structure. Half of the
weight of clay is oxygen, but on a volume basis oxygen is ninety
percent.
The layers of clay are sometimes held together through hydrogen bonds
or potassium bridges and as a result will swell less in the presence of
water. Other clays, such as
montmorillonite, have layers that are loosely attached and will swell greatly when water intervenes between the layers.
There are three groups of clays:
- Crystalline alumino-silica clays: montmorillonite, illite, vermiculite, chlorite, kaolinite.
- Amorphous clays: young mixtures of silica (SiO2-OH) and alumina (Al(OH)3) which have not had time to form regular crystals.
- Sesquioxide clays: old, highly leached clays which result in oxides of iron, aluminium and titanium.
Alumino-silica clays
- Alumino-silica clays are characterised by their regular
crystalline structure. Oxygen in ionic bonds with silicon forms a
tetrahedral coordination (silicon at the center) which in turn forms
sheets of silica. Two sheets of silica are bonded together by a plane of
aluminium which forms an octahedral coordination, called alumina, with
the oxygens of the silica sheet above and that below it. Hydroxyl ions
(OH-) sometimes substitute for oxygen. During the clay formation process, Al3+ may substitute for Si4+, and as much as one fourth of the aluminium Al3+ may be substituted by Zn2+, Mg2+ or Fe2+.
The substitution of lower-valence cations for higher-valence cations
(isomorphic substitution) gives clay a net negative charge that attracts
and holds soil cations, some of which are of value for plant growth.
Isomorphic substitution occurs during the clay's formation and does not
change with time.
-
- Montmorillonite clay is made of four planes of oxygen with
two silicon and one central aluminium plane intervening. The
alumino-silicate montmorillonite clay is said to have a 2:1 ratio of
silicon to aluminium. The seven planes together form a single layer of
montmorillonite. The layers are weakly held together and water may
intervene, causing the clay to swell up to ten times its dry volume. It
occurs in soils which have had little leaching, hence it is found in
arid regions. The entire surface is exposed and available for surface
reactions and it has a high cation exchange capacity (CEC).
-
- 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 and its primary minerals. It has
moderate CEC.
-
- Vermiculite is a mica-based clay similar to illite, but the
layers of clay are held together more loosely by hydrated magnesium and
it will swell, but not as much as does montmorillonite. It has very high
CEC.
-
- Chlorite is similar to vermiculite, but the loose bonding by
occasional hydrated magnesium, as in vermiculite, is replaced by a
hydrated magnesium sheet, that firmly bonds the planes above and below
it. It has two planes of silicon, one of aluminium and one of magnesium;
hence it is a 2:2 clay. Chlorite does not swell and it has low CEC.
-
- 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 replaced 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
specific surface area, and as almost no isomorphic substitution has
occurred it has a low CEC. Where rainfall is high, acid soils
selectively leach more silica than alumina from the original clays,
leaving kaolinite. Even heavier weathering results in sesquioxide clays.
silica-sesquioxide
Amorphous clays
- Amorphous clays are young, and commonly found in volcanic
ash. They are mixtures of alumina and silica which have not formed the
ordered crystal shape of alumino-silica clays which time would provide.
The majority of their negative charges originates from hydroxyl ions,
which can gain or lose a hydrogen ion (H+) in response to
soil pH, in such way was as to buffer the soil pH. They may have either a
negative charge provided by the attached hydroxyl ion (OH-),
which can attract a cation, or lose the hydrogen of the hydroxyl to
solution and display a positive charge which can attract anions. As a
result they may display either high CEC, in an acid soil solution, or
high anion exchange capacity, in a basic soil solution.
Sesquioxide clays
- 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 aluminium hydroxides (Al(OH)3). It takes hundreds of thousands of years of leaching to create sesquioxide clays. Sesqui
is Latin for "one and one-half": there are three parts oxygen to two
parts iron or aluminium; hence the ratio is one and one-half. They are
hydrated and act as either amorphous or crystalline. They are not sticky
and do not swell, and soils high in them behave much like sand and can
rapidly pass water. They are able to hold large quantities of
phosphates. Sesquioxides have low CEC. Such soils range from yellow to
red in colour. Such clays tend to hold phosphorus tightly rendering them
unavailable for absorption by plants.
Organic colloids
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
mineral soil components, it is temporary. It is composed of the very
stable
lignins (30%) and complex sugars (polyuronides,
30%), proteins (30%), waxes, and fats that are resistant to breakdown
by microbes. Its chemical assay is 60% carbon, 5% nitrogen, some oxygen
and the remainder hydrogen, sulfur, and phosphorus. On a dry weight
basis, the CEC of humus is many times greater than that of clay.
Carbon and terra preta
In the extreme environment of high temperatures and the leaching
caused by the heavy rain of tropical rain forests, the clay and organic
colloids are largely destroyed. The heavy rains wash the
alumino-silicate clays from the soil leaving only sesquioxide clays of
low CEC. The high temperatures and humidity allow bacteria and fungi to
virtually dissolve any organic matter on the rain-forest floor overnight
and much of the nutrients are volatilized or leached from the soil and
lost. However, carbon in the form of charcoal is far more stable than
soil colloids and is capable of performing many of the functions of the
soil colloids of sub-tropical soils. Soil containing substantial
quantities of charcoal, of an anthropogenic origin, is called
terra preta.
Research into terra preta is still young but is promising. Fallow
periods "on the Amazonian Dark Earths can be as short as 6 months,
whereas fallow periods on oxisols are usually 8 to 10 years long"
Soil Chemistry
The chemistry of a soil determines its ability to supply available
plant nutrients
and affects its physical properties and the health of its microbial
population. In addition, a soil's chemistry also determines its
corrosivity, stability, and ability to absorb pollutants and to filter
water. It is the surface chemistry of mineral and organic colloids
that determines soil's chemical properties. “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 called humus as well as the
inorganic colloidal particles of clays.” The very high specific surface area
of colloids and their net charges, gives soil its great ability to hold
and release ions. Negatively charged sites on colloids attract and
release cations in what is referred to as cation exchange. Cation-exchange capacity
(CEC) is the amount of exchangeable cations per unit weight of dry soil
and is expressed in terms of milliequivalents of positively charged
ions per 100 grams of soil (or centimoles of positive charge per
kilogram of soil; cmolc/kg). Similarly, positively charged sites on colloids can attract and release anions in the soil giving the soil anion exchange capacity (AEC).
Cation and anion exchange
The
cation exchange, that takes place 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 or positive charges on colloid particles make them able
to hold cations or anions, respectively, to their surfaces. The charges
result from four sources.
- Isomorphous substitution occurs in clay when lower-valence cations
substitute for higher-valence cations in the crystal structure.
Substitutions in the outermost layers are 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.
- Edge-of-clay oxygen atoms are not in balance ionically as the tetrahedral and octahedral structures are incomplete.
- Hydroxyls may substitute for oxygens of the silica layers. When the
hydrogens of the clay hydroxyls are ionised into solution, they leave
the oxygen with a negative charge.
- Hydrogens of humus hydroxyl groups may be ionised into solution, leaving an oxygen with a negative charge.
Cations held to the negatively charged colloids resist being washed
downward by water and out of reach of plants' roots, thereby preserving
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 by the colloid and hence
their ability to replace one another. If present in equal amounts in the
soil water solution:
Al
3+ replaces H
+ replaces Ca
2+ replaces Mg
2+ replaces K
+ same as NH
4+ replaces Na
+
If one cation is added in large amounts, it may replace the others by
the sheer force of its numbers. This is called mass action. This is
largely what occurs with the addition of fertiliser.
As the soil solution becomes more acidic (an abundance of H
+),
the other cations more weakly bound to colloids are pushed into
solution and hydrogen ions occupy those sites. A low pH may cause
hydrogen of hydroxyl groups to be pulled into solution, leaving charged
sites on the colloid available to be occupied by other cations. This
ionisation of hydroxyl groups on the surface of soil colloids creates
what is described as pH-dependent charges. Unlike permanent charges
developed by isomorphous substitution, pH-dependent charges are variable
and increase with increasing pH.
Freed cations can be made available to plants but are also prone to be
leached from the soil, possibly making the soil less fertile. Plants are
able to excrete H
+ into the soil and by that means, change
the pH of the soil near the root and push cations off the colloids, thus
making those available to the plant.
Cation exchange capacity (CEC)
- Cation exchange capacity should be thought of as the soil's
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. 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 relative sterility of tropical
soils. Live plant roots also have some CEC.
Cation exchange capacity for soils; soil textures; soil colloids
| Soil |
State |
CEC meq/100 g |
| Charlotte fine sand |
Florida |
1.0 |
| Ruston fine sandy loam |
Texas |
1.9 |
| Glouchester loam |
New Jersey |
11.9 |
| Grundy silt loam |
Illinois |
26.3 |
| Gleason clay loam |
California |
31.6 |
| Susquehanna clay loam |
Alabama |
34.3 |
| Davie mucky fine sand |
Florida |
100.8 |
| Sands |
------ |
1 - 5 |
| Fine sandy loams |
------ |
5-10 |
| Loams and silt loams |
----- |
5-15 |
| Clay loams |
----- |
15-30 |
| Clays |
----- |
over 30 |
| Sesquioxides |
----- |
0-3 |
| Kaolinite |
----- |
3-15 |
| Illite |
----- |
25-40 |
| Montmorillonite |
----- |
60-100 |
| Vermiculite (similar to illite) |
----- |
80-150 |
| Humus |
----- |
100-300 |
Anion exchange capacity (AEC)
- Anion exchange capacity should be thought of as the soil's
ability to remove anions from the soil water solution and sequester
those for later exchange as the plant roots release carbonate anions to
the soil water solution. Those colloids which have low CEC tend to have
some AEC. Amorphous and sesquioxide clays have the highest AEC, followed
by the iron oxides. Levels of AEC are 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 SO42- replaces NO3- replaces Cl-
- The amount of exchangeable anions is of a magnitude of tenths to a few milliequivalents per 100 g dry soil.
As pH rises, there are relatively more hydroxyls, which will displace
anions from the colloids and force them into solution and out of
storage; hence AEC decreases with increasing pH (alkalinity).
Soil reaction (pH)
Soil reactivity is expressed in terms of pH and is a measure of the
acidity or alkalinity 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
- At 25°C an aqueous solution that has a pH of 3.5 has 10-3.5 moles H+ (hydrogen ions) per litre of solution (and also 10-10.5 mole/litre OH-). A pH of 7, defined as neutral, has 10−7 moles hydrogen ions per litre of solution and also 10−7 moles of OH- per litre; since the two concentrations are equal, they are said to neutralise each other. A pH of 9.5 has 10-9.5 moles hydrogen ions per litre of solution (and also 10-2.5 mole per litre OH-). A pH of 3.5 has one million times more hydrogen ions per litre than a solution with pH of 9.5 (9.5 - 3.5 = 6 or 106) and is more acidic.
- The effect of pH on a soil is to remove from the soil or to make
available certain ions. Soils with high acidity tend to have toxic
amounts of aluminium and manganese. Plants which 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 with mineral soils of pH 6.5 and organic soils of pH 5.5.
- In high rainfall areas, soils tend to acidity as the basic cations
are forced off the soil colloids by the mass action of hydrogen ions
from the rain as those attach to the colloids. High rainfall rates can
then wash the nutrients out, leaving the soil sterile. Once the colloids
are saturated with H+, the addition of any more hydrogen
ions or aluminum hydroxyl cations drives the pH even lower (more acidic)
as the soil has been left with no buffering capacity. In areas of
extreme rainfall and high temperatures, the clay and humus may be washed
out, further reducing the buffering capacity of the soil. 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 nutrients can supply the deficit. Sodium
can be reduced by the addition of gypsum
(calcium sulphate) as calcium adheres to clay more tightly than does
sodium causing sodium to be pushed into the soil water solution where it
can be washed out by an abundance of water.
Base saturation percentage
- There are acid-forming cations (hydrogen and aluminium) 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 aluminium and
hydrogen cations (acid-forming), the remainder of positions on the
colloids (20-5 = 15 meq) are assumed occupied by base-forming cations,
so that the percentage base saturation is 15/20 x 100% = 75% (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 (increases with increasing pH). It is of use in calculating the
amount of lime needed to neutralise an acid soil. The amount of lime
needed to neutralize a soil must take account of the amount of acid
forming ions on the colloids not just those in the soil water solution.
The addition of enough lime to neutralize the soil water solution will
be insufficient to change the pH, as the acid forming cations stored on
the soil colloids will tend to restore the original pH condition as they
are pushed off those colloids by the calcium of the added lime.
Buffering of soils
The resistance of soil to changes in pH as a result of 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, while soils high in colloids have high buffering
capacity. Buffering occurs by cation exchange and neutralisation.
The addition of a small amount highly basic aqueous ammonia to a soil
will cause the ammonium to displace hydrogen ions from the colloids,
and the end product is water and colloidally fixed ammonium, but no
permanent change overall in soil pH.
The addition of a small amount of
lime, CaCO3, will displace hydrogen ions from the soil colloids, causing the fixation of calcium to colloids and the evolution of CO2 and water, with no permanent change in soil pH.
The addition of carbonic acid (the solution of CO
2 in
water) will displace calcium from colloids, as hydrogen ions are fixed
to the colloids, evolving water and slightly alkaline (temporary
increase in pH) highly soluble calcium bicarbonate, which will then
precipitate as lime (CaCO
3) and water at a lower level in the soil profile, with the result of no permanent change in soil pH.
All of the above are examples of the buffering of soil 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
Sixteen nutrients are 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.
Nutrients required for plants to complete their life cycle are
considered essential nutrients. Nutrients that enhance the growth of
plants but are not necessary to complete the plant's life cycle are
considered non-essential. With the exception of carbon, hydrogen and
oxygen, which are supplied by carbon dioxide and water, the nutrients
derive originally from the mineral component of the soil. Although
minerals are the origin of those nutrients, the organic component of the
soil is the reservoir of the majority of readily available plant
nutrients. For the nutrients to be available to plants, they must be in
the proper ionic form (with the exception of water and CO2). For example, the application of finely ground minerals, feldspar and apatite, to soil does not provide the necessary amounts of potassium and phosphorus for good plant growth.
Nitrogen is the primary limiting nutrient and phosphorus is second to
nitrogen as the primary nutrient for plants, animals and microorganisms.
The provision of plant nutrition involves chemical, biological, and
physical processes. Nearly all plant nutrients are taken up from the
soil water solution in the form of
ions, either cations or as anions. In an effort to gain nutrients, plants will release ions to the soil. Bicarbonate (HCO3-) and hydroxyl (OH-)
anions released from plant roots enhance the absorption of nutrient
anions; similarly, hydrogen cations are released in exchange for cation
forms of nutrients. As a result, nutrient ions are pushed into the soil
water solution from their sequestration on colloids to become available
to plants. Nitrogen, for example, is available in soil organic material
but is unusable by plants until it is made available by that material's
decomposition by micro-organisms into cation or anion forms. The NH4+ (ammonium) and NO3-
(nitrate) forms of nitrogen are stored on the soil colloids until
forced off those by the presence of other cations and anions. After
that, they will move by physical means to near the plant roots.
Generally, plant roots can readily absorb all of the nutrients from the
soil solution, provided there is enough oxygen gas in the soil to
support root metabolism.
The bulk of most nutrient elements in the soil is held in the form of
primary and secondary minerals, and organic matter. The primary
minerals (mostly rock dust in the form of silt) hold the nutrients too
tightly to be readily available; the nutrients adsorbed onto the
colloids clay and humus are moderately available; and the soil solution
fraction has ions that are freely available for absorption by plant
roots. Gram for gram, the capacity of humus to hold nutrients and water
is far greater than that of clay. All in all, small amounts of humus may
remarkably increase the soil’s capacity to promote plant growth.
Plant nutrients, their chemical symbols, and the ionic forms common in soils and available for plant uptake
| Element |
Symbol |
Ion or molecule |
| Carbon |
C |
CO2 (mostly through leaves) |
| Hydrogen |
H |
H+, HOH (water) |
| Oxygen |
O |
O2-, OH -, CO32-, SO42-, CO2 |
| Phosphorus |
P |
H2PO4 -, HPO42- (phosphates) |
| Potassium |
K |
K+ |
| Nitrogen |
N |
NH4+, NO3 - (ammonium, nitrate) |
| Sulfur |
S |
SO42- |
| Calcium |
Ca |
Ca2+ |
| Iron |
Fe |
Fe2+, Fe3+ (ferrous, ferric) |
| Magnesium |
Mg |
Mg2+ |
| Boron |
B |
H3BO3, H2BO3 -, B(OH)4 - |
| Manganese |
Mn |
Mn2+ |
| Copper |
Cu |
Cu2+ |
| Zinc |
Zn |
Zn2+ |
| Molybdenum |
Mo |
MoO42- (molybdate) |
| Chlorine |
Cl |
Cl - (chloride) |
Mechanism of nutrient uptake
All the nutrients with the exception of carbon are taken up by the
plant through its roots. To be taken up by a plant, a nutrient element
must be in an ionic form (with the exception of water and H
3BO
3)
and must be located at the root surface. Often, parts of a root are in
such intimate contact with soil particles that a direct exchange may
take place between nutrient ions adsorbed on the surface of the soil
colloids and hydrogen ions from the surface of root cell walls. In any
case, the supply of nutrients in contact with the root would soon be
depleted. There are however, three basic mechanisms by which the
concentration of nutrient ions at the root surface is maintained.
Nutrient levels are maintained by three principal mechanisms by which
nutrient ions dissolved in the soil solution are brought into contact
with plant roots:
- Mass flow
- Diffusion
- Root interception
All three mechanisms operate simultaneously, but one mechanism or
another may be most important for a particular nutrient. For example, in
the case of calcium, which is generally plentiful in the soil solution,
mass flow alone can usually bring sufficient amounts to the root
surface. However, in the case of phosphorus, diffusion is needed to
supplement mass flow because the soil solution is very low in this
element in comparison to the amounts needed by plants. For the most
part, nutrient ions must travel some distance in the soil solution to
reach the root surface. This movement can take place by
mass flow,
as when dissolved nutrients are carried along with the soil water
flowing toward a root that is actively drawing water from the soil. In
this type of movement, the nutrient ions are somewhat analogous to
leaves floating down a stream. In addition, nutrient ions continually
move by
diffusion from areas of greater concentration toward the
nutrient-depleted areas of lower concentration around the root surface.
By this means, plants can continue to take up nutrients even at night,
when water is only slowly absorbed into the roots. Finally,
root interception comes into play as roots continually grow into new, undepleted soil.
Because nutrient uptake is an active metabolic process, conditions
that inhibit root metabolism may also inhibit nutrient uptake. Examples
of such conditions include excessive soil water content or soil
compaction resulting in poor soil aeration, excessively high or low soil
temperatures, and above-ground conditions that result in low
translocation of sugars to plant roots. A
maize plant will use one quart of water per day at the height of its growing season.
Estimated relative importance of mass flow, diffusion and
root interception as mechanisms in supplying plant nutrients to corn
plant roots in soils
| Nutrient |
Approximate percentage supplied by: |
| Mass flow |
Root interception |
Diffusion |
| Nitrogen |
98.8 |
1.2 |
0 |
| Phosphorus |
6.3 |
2.8 |
90.9 |
| Potassium |
20.0 |
2.3 |
77.7 |
| Calcium |
71.4 |
28.6 |
0 |
| Sulfur |
95.0 |
5.0 |
0 |
| Molybdenum |
95.2 |
4.8 |
0 |
In the above table, phosphorus and potassium nutrients move more by
diffusion than they do by mass flow in water solution, as they are
rapidly taken up by the roots creating a concentration of almost zero
near the roots. The very steep concentration gradient is of greater
influence in the movement of those ions than is the movement of those by
mass flow.
The movement by mass flow requires the transpiration of water from the
plant causing water and solution ions to also move toward the roots.
Movement by root interception is slowest as the plants must extend their
roots. Plants move ions out of their roots in an effort to move
nutrients in from the soil. Hydrogen H
+ is exchanged for cations, and carbonate (HCO
3-) and hydroxide (OH
-)
anions are exchanged for nutrient anions. Plants derive most of their
anion nutrients from decomposing organic matter, which holds 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 to the soil water as other ions move off of clay and
humus, are added from the decomposition of soil minerals, and are
released by the decomposition of organic matter. Where crops are
produced, the replenishment of nutrients in the soil must be augmented
by the addition of fertiliser or organic matter.
Carbon
Measuring soil respiration in the field using an SRS2000 system.
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
metabolised and carbon dioxide (CO
2) is released as a
byproduct which then finds its way out of the soil and into the
atmosphere. The nitrogen, and other nutrients however, is sequestered in
the bodies of the living matter of those organisms and so it builds up
in the soil. Normal CO
2 concentration in the atmosphere is
0.03%, which is probably the factor limiting plant growth. In a field of
maize on a still day during high light conditions in the growing
season, the CO
2 concentration drops very low, but under such
conditions the crop could use up to 20 times the normal concentration.
The respiration of CO
2 by soil micro-organisms decomposing soil organic matter contributes an important amount of CO
2 to the photosynthesising plants. Within the soil, CO
2
concentration is 10 to 100 times that of atmospheric levels but may
rise to toxic levels if the soil porosity is low or if diffusion is
impeded by flooding.
Nitrogen
Generalization of percent soil nitrogen by soil order
Nitrogen is the most critical element obtained by plants from the soil and is a bottleneck in plant growth
. Plants can use the nitrogen as either the ammonium cation (NH
4+) or the anion nitrate (NO
3-).
Nitrogen is seldom missing in the soil, but is often in the form of raw
organic material which cannot be used directly. The total nitrogen
content depends on the climate, vegetation, topography, age and soil
management. Soil nitrogen typically decreases by 0.2 to 0.3% for every
temperature increase by 10 °C. Usually, more nitrogen is under grassland
than under forest. Humus formation promotes nitrogen immobilization.
Cultivation decreases soil nitrogen by exposing soil to more air which
the bacteria can use, and no-tillage maintains more nitrogen than
tillage.
Carbon/Nitrogen Ratio of Various Organic Materials
Sortable table
| Alfalfa |
13 |
| Bacteria |
4 |
| Clover, green sweet |
16 |
| Clover, mature sweet |
23 |
| Fungi |
9 |
| Forest litter |
30 |
| Humus in warm cultivated soils |
11 |
| Legume-grass hay |
25 |
| Legumes (alfalfa or clover), mature |
20 |
| Manure, cow |
18 |
| Manure, horse |
16–45 |
| Manure, human |
10 |
| Oat straw |
80 |
| Straw, cornstalks |
90 |
| Sawdust |
250 |
Some micro-organisms are able to metabolise organic matter and release ammonium in a process called
mineralisation. Others take free ammonium and oxidise it to nitrate. Particular bacteria are capable of metabolising N
2 into the form of nitrate in a process called
nitrogen fixation. Both ammonium and nitrate can be immobilized
or essentially lost from the soil by its incorporation into the
microbes' living cells, where it is temporarily sequestered in the form
of amino acids and protein. Nitrate may also be lost from the soil when
bacteria metabolise it to the gases N2 and N2O. The loss of gaseous forms of nitrogen to the atmosphere due to microbial action is called denitrification. Nitrogen may also be leached
from the soil if it is in the form of nitrate or lost to the atmosphere
as ammonia due to a chemical reaction of ammonium with alkaline soil by
way of a process called volatilisation. Ammonium may also be sequestered in clay by fixation. A small amount of nitrogen is added to soil by rainfall.
Nitrogen gains
In the process of
mineralisation, microbes 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 the immobilization process. In that form the nitrogen is said to be immobilised. Later, when such bacteria die, they too are mineralised
and some of the nitrogen is released as ammonium and nitrate. If the
C/N is less than 15, ammonia is freed to the soil, where it may be used
by bacteria which oxidise it to nitrate (nitrification).
Bacteria may on average add 25 pounds (11 kg) nitrogen per acre, and in
an unfertilised field, this is the most important source of usable
nitrogen. In a soil with 5% organic matter perhaps 2 to 5% of that is
released to the soil by such decomposition. It occurs fastest in warm,
moist, well aerated soil. The mineralisation of 3% of the organic
material of a soil that is 4% organic matter overall, would release 120
pounds (54 kg) of nitrogen as ammonium per acre.
In
nitrogen fixation, rhizobium bacteria convert N2 to nitrate (NO3-). Rhizobia share a symbiotic relationship
with host plants, since rhizobia supply the host with nitrogen and the
host provides rhizobia 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 mineralisation.
Some amount of usable nitrogen is fixed by
lightning as nitric oxide (NO) and nitrogen dioxide (NO2-). Nitrogen dioxide is soluble in water to form nitric acid (HNO3) solution of H+ and NO3-. Ammonia, NH3,
previously released from the soil or from combustion, may fall with
precipitation as nitric acid at a rate of about five pounds nitrogen per
acre per year.
Nitrogen sequestration
When bacteria feed on soluble forms of nitrogen (ammonium and
nitrate), they temporarily sequester that nitrogen in their bodies in a
process called
immobilisation. At a later time when those bacteria die, their nitrogen may be released as ammonium by the processes of mineralisation.
Protein material is easily broken down, but the rate of its
decomposition is slowed by its attachment to the crystalline structure
of clay and when trapped 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 pushes
potassium ions from between the layers of clay such as illite or montmorillonite. Only a small fraction of nitrogen is held this way.
Nitrogen losses
Usable nitrogen may be lost from soils when it is in the form of nitrate, as it is easily
leached. Further losses of nitrogen occur by denitrification, the process whereby soil bacteria convert nitrate (NO3-) to nitrogen gas, N2 or N2O. This occurs when poor soil aeration
limits free oxygen, forcing bacteria to use the oxygen in nitrate for
their respiratory process. Denitrification increases when oxidisable
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. Denitrification may cause the loss of 10 to
20 percent of the available nitrates within a day and when conditions
are favourable to that process, losses of up to 60 percent of nitrate
applied as fertiliser may occur.
Ammonium volatilisation occurs when ammonium reacts chemically with an alkaline soil, converting NH
4+ to NH
3. The application of ammonium fertiliser to such a field can result in volatilisation losses of as much as 30 percent.
Phosphorus
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 lb of phosphorus per acre in the soil, it is generally
unavailable 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 mineralisation of organic matter. Agricultural fields may
need to be fertilised to make up for the phosphorus that has been
removed in the crop.
When phosphorus does form solubilised ions of H
2PO
4-,
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 fertilisers to soils may result in zinc
deficiencies as zinc phosphates form. Conversely, the application of
zinc to soils may immobilise phosphorus again 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.
Potassium
The amount of potassium in a soil may be as much as 80,000 lb per
acre-foot, of which only 150 lb is available for plant growth. Common
mineral sources of potassium are the mica
biotite and potassium feldspar, KAlSi3O8. When solubilised, half will be held as exchangeable cations on clay while the other half is in the soil water solution. Potassium fixation
often occurs when soils dry and the potassium is bonded between layers
of illite 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.
Potassium may be leached from soils low in clay.
Calcium
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 thus low 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, CaCO
3, is more soluble than potassium-bearing minerals.
Magnesium
Magnesium is central to
chlorophyll
and aids in the uptake of phosphorus. The minimum amount of magnesium
required for plant health is not sufficient for the health of forage
animals. A common mineral source of magnesium is the black mica mineral,
biotite. Magnesium is generally available in soil, but is missing from
some along the Gulf and Atlantic coasts of the United States due to
leaching by heavy precipitation.
Sulfur
Sulfur is essential to the formation of proteins and chlorophyll, and
essential to plant vitamin synthesis. Most sulfur is made available to
plants, like phosphorus, by its release from decomposing organic matter.
Deficiencies may exist in some soils and if cropped, sulfur needs to be
added. The application of large quantities of nitrogen to fields that
have marginal amounts of sulfur may cause sulfur deficiency in the
rapidly growing plants by the plant's growth outpacing the supply of
sulfur. A 15-ton crop of onions uses up to 19 lb of sulfur and 4 tons of
alfalfa uses 15 lb 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 lb per acre respectively.
Micronutrients
The micronutrients essential for plant life, in their order of importance, include
iron, manganese, zinc, copper, boron, chlorine and molybdenum.
The term refers to plants' needs, not to their abundance in soil. They
are required in very small amounts but are essential to plant health in
that most are required parts of some enzyme system which speeds up
plants' metabolisms. 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 zinc and
iron phosphates. Iron deficiency may also result from excessive amounts
of heavy metals or calcium minerals (lime) in the soil. Excess amounts
of soluble boron, molybdenum and chloride are toxic.
Non-essential nutrients
Nutrients which enhance the health but whose deficiency does not stop the life cycle of plants include:
cobalt, strontium, vanadium, silicon and nickel.[citation needed] As their importance are evaluated they may be added to the list of essential plant nutrients.
Soil organic matter
Soil organic matter is made up of
organic compounds
and includes plant, animal and microbial material, both living and
dead. A typical soil has a biomass composition of 70% microorganisms,
22% macrofauna, and 8% roots. The living component of an acre of soil
may include 900 lb of earthworms, 2400 lb of fungi, 1500 lb of bacteria,
133 lb of protozoa and 890 lb of arthropods and algae.
A small part of the organic matter consists of the living cells such
as bacteria, molds, and actinomycetes that work to break down the dead
organic matter. Were it not for the action of these micro-organisms, the
entire carbon dioxide part of the atmosphere would be sequestered as
organic matter in the soil.
Chemically, organic matter is classed as follows:
- 1. Polysaccharides
- a. cellulose
- b. hemicellulose
- c. starch
- d. pectin
- 2. Lignins
- 3. Proteins
Most living things in soils, including plants, insects, bacteria, and
fungi, are dependent on organic matter for nutrients and energy. Soils
have organic compounds in varying degrees of decomposition which rate is
dependent on the temperature, soil moisture, and aeration. Bacteria and
fungi feed on the raw organic matter, which are fed upon by amoebas,
which in turn are fed upon by nematodes and arthropods. Organic matter
holds soils open, allowing the infiltration of air and water, and may
hold as much as twice its weight in water. Many soils, including desert
and rocky-gravel soils, have little or no organic matter. Soils that are
all organic matter, such as
peat (histosols), are infertile.
In its earliest stage of decomposition, the original organic material
is often called raw organic matter. The final stage of decomposition is
called humus.
In grassland, much of the organic matter added to the soil is from
the deep, fibrous, grass root systems. By contrast, tree leaves falling
on the forest floor are the principal source of soil organic matter in
the forest. Another difference is the frequent occurrence in the
grasslands of fires that destroy large amounts of aboveground material
but stimulate even greater contributions from roots. Also, the much
greater acidity under any forests inhibits the action of certain soil
organisms that otherwise would mix much of the surface litter into the
mineral soil. As a result, the soils under grasslands generally develop a
thicker A horizon with a deeper distribution of organic matter than in
comparable soils under forests, which characteristically store most of
their organic matter in the forest floor (O horizon) and thin A horizon.
Humus
Humus
refers to organic matter that has been decomposed by soil flora and
fauna to the 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 crucial to soil health
and plant growth. Humus also hold bits of undecomposed organic matter
which feed arthropods and worms which further improve the soil. The end
product, humus, is soluble in water and forms a weak acid that can
attack silicate minerals.
Humus has a high cation exchange capacity that on a dry weight basis is
many times greater than that of clay colloids. It also acts as a
buffer, like clay, against changes in pH and soil moisture.
Humic acids and fulvic acids,
which begin as raw organic matter, are important constituents of humus.
After the death of plants and animals, microbes begin to feed on the
residues, resulting finally in the formation of humus. With
decomposition, there is a reduction of water-soluble constituents, cellulose and hemicellulose,
and nutrients such as nitrogen, phosphorus, and sulfur. As the residues
break down, only stable molecules made of aromatic carbon rings, oxygen
and hydrogen remain in the form of humin, lignin
and lignin complexes collectively called humus. While the structure of
humus has few nutrients, it is able to attract and hold cation and anion
nutrients by weak bonds that can be released into the soil solution in
response to changes in soil pH.
Lignin is resistant to breakdown and accumulates within the soil. It
also reacts with amino acids, which further increases 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 and are stabilised by the clay. Proteins normally decompose readily,
but when bound to clay particles, they become more resistant to
decomposition. Clay particles also absorb the enzymes exuded by microbes
which 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. High soil tannin (polyphenol) content can cause nitrogen to be sequestered in proteins or cause nitrogen immobilisation.
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 typically have 3 to 6 percent nitrogen. Raw
organic matter, as a reserve of nitrogen and phosphorus, is a vital
component affecting
soil fertility. Humus also absorbs water, and expands and shrinks between dry and wet states, increasing soil porosity.
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 organic matter
The production, accumulation and degradation of organic matter are greatly dependent on climate. Temperature, soil moisture and
topography
are the major factors affecting the accumulation of organic matter in
soils. Organic matter tends to accumulate under wet or cold conditions
where decomposer activity is impeded by low temperatur or excess moisture which results in anaerobic conditions.
Conversely, excessive rain and high temperatures of tropical climates
enables rapid decomposition of organic matter and leaching of plant
nutrients; forest ecosystems on these soils rely on efficient recycling
of nutrients and plant matter to maintain their productivity.
Excessive slope may encourage the erosion of the top layer of soil
which holds most of the raw organic material that would otherwise
eventually become humus.
Plant residue in soil
Typical types and percentages of plant residue components
Cellulose (45%)
Lignin (20%)
Hemicellulose (18%)
Protein (8%)
Sugars and starches (5%)
Fats and waxes (2%)
Cellulose and hemicellulose undergo fast decomposition by fungi and bacteria, with a half-life of 12–18 days in a temperate climate. Brown rot fungi can decompose the cellulose and hemicellulose, leaving the lignin and phenolic compounds behind. Starch,
which is an energy storage system for plants, undergoes fast
decomposition by bacteria and fungi. Lignin consists of polymers
composed of 500 to 600 units with a highly branched, amorphous
structure. Lignin undergoes very slow decomposition, mainly by white rot fungi and actinomycetes; its half-life under temperate conditions is about six months.
Soil horizons
Main article:
Soil horizon
A horizontal layer of the soil, whose physical features, composition
and age are distinct from those above and beneath, are referred to as a
soil horizon.
The naming of a horizon is based on the type of material of which it is
composed. Those materials reflect the duration of specific processes of
soil formation. They are labelled using a shorthand notation of letters
and numbers
which describe the horizon in terms of its colour, size, texture,
structure, consistency, root quantity, pH, voids, boundary
characteristics and presence of nodules or concretions. Few soil profiles have all the major horizons. Some may have only one horizon.
The exposure of parent material to favourable conditions produces
mineral soils that are marginally suitable for plant growth. That 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. Organisms colonise 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
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 which 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) aims to establish an international reference base for soil classification.
Soil classification systems
Australia
There are fourteen soil orders at the top level of the Australian
Soil Classification. They are: Anthroposols, Organosols, Podosols,
Vertosols, Hydrosols, Kurosols, Sodosols, Chromosols, Calcarosols,
Ferrosols, Dermosols, Kandosols, Rudosols and Tenosols.
European Union
The EU's soil taxonomy is based on a new standard soil classification
in the World Reference Base for Soil Resources produced by the
UN's Food and Agriculture Organization. According to this, the major soils in the European Union are:
USA
A taxonomy is an arrangement in a systematic manner; the
USDA soil taxonomy has six levels of classification. They are, from most general to specific: order, suborder, great group, subgroup, family and series.
Soil properties that can be measured quantitatively are used in this
classification system — they include: depth, moisture, temperature,
texture, structure, cation exchange capacity, base saturation, clay
mineralogy, organic matter content and salt content. There are 12 soil
orders (the top hierarchical level) in the USDA soil taxonomy. The names of the orders end with the suffix -sol. The criteria for the different soil orders include properties that reflect major differences in the genesis of soils. The orders are:
- 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
constitute 10.1% of soils worldwide.
- Andisols - volcanic ash soils. They are young and very fertile. They cover 1% of the world's ice-free surface.
- Aridisol - dry soils forming under desert
conditions which have fewer than 90 consecutive days of moisture during
the growing season and are nonleached. 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 metres of the surface or
gelic materials and permafrost within one metre. They constitute 9.1% of
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 constitute 15% of soils worldwide.
- Mollisols
- 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 heavily weathered, are rich in iron and aluminum oxides (sesquioxides) or kaolin but low in silica. They have only trace nutrients due to heavy tropical rainfall and high temperatures. They are 7.5% of soils worldwide.
- Spodosol
- acid soils with organic colloid layer complexed with iron and
aluminium leached from a layer above. They are typical soils of coniferous and deciduous forests in cooler climates. They constitute 4% of soils worldwide.
- Ultisol - acid soils in the humid tropics and subtropics, which are depleted in calcium, magnesium and potassium (important plant nutrients). They are highly weathered, but not as weathered as Oxisols. They make up 8% 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 into which surface layers
can fall. They are difficult to farm or to construct roads and buildings
due to their high expansion rate. They constitute 2.4% of soils
worldwide.
The percentages listed above
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" which have no soil.
The above 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 which 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 in which the kind and sequence
of soil horizons distinguish one soil from another. About 185 great
groups are recognized in the United States. Horizons marked by clay,
iron, humus and hard pans and soil features such as the
expansion-contraction of clays (that produce self-mixing provided by
clay), temperature, and marked quantities of various salts are used as
distinguishing features.
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 towards (is similar to) soils of other soil great groups,
suborders or orders. These properties are not developed or expressed
well enough to cause the soil to be included within the great group
towards which they grade, but suggest similarities. Extragrade features
are aberrant properties which prevent that soil from being included in
another soil classification. About 1,000 soil subgroups are defined in
the United States.
A
soil family category is a group of soils within a subgroup
and describes the physical and chemical properties which affect the
response of soil to agricultural management and engineering
applications. The principal characteristics used to differentiate soil
families include texture, mineralogy, pH, permeability, structure,
consistency, the locale'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 recognised in the United States.
A family may contain several
soil series which describe the
physical location using the name of a prominent physical feature such as
a river or town near where the soil sample was taken. An example would
be Merrimac for the
Merrimack River
in New Hampshire, USA. More than 14,000 soil series are recognised in
the United States. This permits very specific descriptions of soils.
A
soil phase of series, originally called 'soil type' describes the soil surface texture, slope, stoniness, saltiness, erosion, and other conditions.
Uses
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 can
be dissolved in a solution. The types of soil and available moisture
determine the species of plants that can be cultivated.
Soil material is also a critical component in the mining, construction and landscape development 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 fibre production. Soil provides minerals and water to plants. Soil
absorbs rainwater and releases it later, thus preventing floods and
drought. Soil cleans 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) below-ground. Above-ground and below-ground biodiversities are tightly interconnected, 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.
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 waste water 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.
Geophagy
is the practice of eating soil-like substances. Both animals and human
cultures occasionally consume soil for medicinal, recreational, or
religious purposes. It has been shown that some monkeys consume soil, together with their preferred food (tree foliage and fruits), in order to alleviate tannin toxicity.
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. Soil organisms metabolise them or immobilise them in their biomass and necromass, thereby incorporating them into stable humus. The physical integrity of soil is also a prerequisite for avoiding landslides in rugged landscapes.
Degradation
Land degradation refers to a 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.
While
soil acidification
is beneficial in the case of alkaline soils, it degrades land when it
lowers crop productivity and increases soil vulnerability to
contamination and erosion. Soils are often initially acid because their parent materials were acid and initially low in the basic cations (calcium, magnesium, potassium and sodium).
Acidification occurs when these elements are leached from the soil
profile by rainfall or the by harvesting of forest or agricultural
crops. Soil acidification is accelerated by the use of acid-forming nitrogenous fertilizers and by the effects of acid precipitation.
Soil contamination
at low levels is often within soil's capacity to treat and assimilate
waste material. Soil biota can treat waste by transforming it; soil
colloids can adsorb waste material. Many waste treatment processes rely
on this treatment capacity. Exceeding treatment capacity can damage soil
biota and limit soil function. Derelict soils occur where industrial
contamination or other development activity damages the soil to such a
degree that the land cannot be used safely or productively. Remediation of derelict soil uses principles of geology, physics, chemistry and biology to degrade, attenuate, isolate or remove soil contaminants to restore soil functions and values. Techniques include leaching, air sparging, chemical amendments, phytoremediation, bioremediation and natural degradation.
Desertification
is an environmental process of ecosystem degradation in arid and
semi-arid regions, often caused by human activity. It is a common
misconception that droughts
cause desertification. Droughts are common in arid and semiarid lands.
Well-managed lands can recover from drought when the rains return. Soil
management tools include maintaining soil nutrient and organic matter
levels, reduced tillage and increased cover. These practices help to
control erosion and maintain productivity during periods when moisture
is available. Continued land abuse during droughts, however, increases
land degradation. Increased population and livestock pressure on
marginal lands accelerates desertification.
Erosion of soil is caused by water, wind, ice, and movement in response to gravity. More than one kind of erosion can occur simultaneously. Erosion is distinguished from weathering, since erosion also transports eroded soil away from its place of origin (soil in transit may be described as sediment). Erosion is an intrinsic natural process, but in many places it is greatly increased by human activity, especially poor land use practices. These include agricultural activities which leave the soil bare during times of heavy rain or strong winds, overgrazing, deforestation, and improper construction activity. Improved management can limit erosion. Soil conservation techniques which are employed include changes of land use (such as replacing erosion-prone crops with grass or other soil-binding plants), changes to the timing or type of agricultural operations, terrace building, use of erosion-suppressing cover materials (including cover crops and other plants), limiting disturbance during construction, and avoiding construction during erosion-prone periods.
A serious and long-running water erosion problem occurs in
China, on the middle reaches of the Yellow River and the upper reaches of the Yangtze River. From the Yellow River, over 1.6 billion tons of sediment flow each year into the ocean. The sediment originates primarily from water erosion (gully erosion) in the Loess Plateau region of northwest China.
Soil piping is a particular form of soil erosion that occurs below
the soil surface. It causes levee and dam failure, as well as
sink hole formation. Turbulent flow removes soil starting at the mouth of the seep flow and the subsoil erosion advances up-gradient. The term sand boil is used to describe the appearance of the discharging end of an active soil pipe.
Soil salination is the accumulation of free salts
to such an extent that it leads to degradation of the agricultural
value of soils and vegetation. Consequences include corrosion damage,
reduced plant growth, erosion due to loss of plant cover and soil
structure, and water quality problems due to sedimentation. Salination
occurs due to a combination of natural and human-caused processes. Arid
conditions favour salt accumulation. This is especially apparent when
soil parent material is saline. Irrigation of arid lands is especially problematic.
All irrigation water has some level of salinity. Irrigation, especially
when it involves leakage from canals and overirrigation in the field,
often raises the underlying water table. Rapid salination occurs when the land surface is within the capillary fringe of saline groundwater. Soil salinity control involves watertable control and flushing with higher levels of applied water in combination with tile drainage or another form of subsurface drainage.
Reclamation
Soils which contain high levels of particular clays, such as
smectites, are often very fertile. For example, the smectite-rich clays of Thailand's Central Plains are among the most productive in the world.
Many farmers in tropical areas, however, struggle to retain organic
matter in the soils they work. In recent years, for example,
productivity has declined in the low-clay soils of northern Thailand.
Farmers initially responded by adding organic matter from termite
mounds, but this was unsustainable in the long-term. Scientists
experimented with adding
bentonite, one of the smectite family of clays, to the soil. In field trials, conducted by scientists from the International Water Management Institute in cooperation with Khon Kaen University
and local farmers, this had the effect of helping retain water and
nutrients. Supplementing the farmer's usual practice with a single
application of 200 kg bentonite per rai (6.26 rai = 1 hectare) resulted
in an average yield increase of 73%. More work showed that applying
bentonite to degraded sandy soils reduced the risk of crop failure
during drought years.
In 2008, three years after the initial trials,
IWMI
scientists conducted a survey among 250 farmers in northeast Thailand,
half of whom had applied bentonite to their fields. The average
improvement for those using the clay addition was 18% higher than for
non-clay users. Using the clay had enabled some farmers to switch to
growing vegetables, which need more fertile soil. This helped to
increase their income. The researchers estimated that 200 farmers in
northeast Thailand and 400 in Cambodia had adopted the use of clays, and
that a further 20,000 farmers were introduced to the new technique.
If the soil is too high in clay, adding gypsum, washed river sand and
organic matter will balance the composition. Adding organic matter
(like
ramial chipped wood for instance) to soil which is depleted in nutrients and too high in sand will boost its quality.
