Scope of Soil
Science
Soil
Science has six well defined and developed disciplines. Scope of soil Science
is reflected through these disciplines.
Soil Science : The science dealing with soil as
a natural resource on the surface of the earth, including Pedology (soil
genesis, classification and mapping) and the physical, chemical and biological
and fertility properties of soil and these properties in relation to their
management for crop production.
1.
Soil fertility:
Nutrient supplying properties of soil
2.
Soil chemistry:
Chemical constituents, chemical properties and the chemical reactions
3.
Soil physics:
Involves the study of physical properties
4.
Soil microbiology: deals with micro organisms, its population, classification, its
role in transformations
5.
Soil conservation: Dealing with protection of soil against physical loss by erosion
or against chemical deterioration i.e. excessive loss of nutrients either
natural or artificial means.
6.
Pedology: Dealing with
the genesis, survey and classification
Soil can be compared to various systems of human body
Digestive
- matters decomposition
Respiratory - air circulation & exchange of gases
Circulatory - water movement with in the soil system
Excretory - leaching out of excess salts
Brain - soil clay
Colour - soil colour
Height - soil depth
Components of Soil (Volume basis)
Mineral
matter – 45%
Organic matter – 5%
Soil water – 25%
Soil air – 25%
Definition
of Soil & Approaches of Soil Study
Definition of Soil
Whitney (1982) Hilgard (1892) Dokuchaiev (1900)
Joffe (1936): Soil is a natural body of mineral and organic constituents
differentiated into horizons usually unconsolidated, of variable depth which
differs among themselves as well as from the underlying parent material in
morphology, physical makeup, chemical properties and composition and biological
characteristics.
SSSA (1970):
(i) The unconsolidated mineral matter on the surface of the earth that has been
subjected to and influenced by genetic and environmental factors of parent
material, climate (including moisture and temperature effects), macro and
microorganisms and topography, all affecting over a period of time and
producing a product, that is “SOIL” that differs from the material from which
it is derived in many, physical, chemical, biological and morphological
properties and characteristics.
(ii) The unconsolidated mineral material on the
immediate surface of the earth that serves as a natural medium for the growth
of land plants.
Approaches of Soil Study
Two Concepts: One treats soil as a natural body,
weathered and synthesized product in nature while other treats soil as a medium
for plant growth.
1) Pedological Approach: The origin of the soil, its classification and
its description are examined in Pedology. (From Greek word pedon, means
soil or earth). Pedology is the study of soil as a natural body and does not
focus on the soil’s immediate practical use. A pedologist studies, examines and
classifies soil as they occur in their natural environment.
2) Edaphological
Approach: Edophology (from
Greek word edaphos, means soil or ground) is the study of soil from the
stand point of higher plants. Edaphologists consider the various properties of
soil in relation to plant production. They are practical and have the
production of food and fiber as their ultimate goal. They must determine the
reasons for variation in the productivity of soils and find means for
improvement.
Weathering
of Rocks and Minerals
Rocks and minerals are formed under a very high
temperature and pressure, exposed to atmospheric conditions of low pressure and
low temperature and they become unstable and weather.
Soils are formed from rocks through the intermediate
stage of formation of Regolith which is the resultant of weathering.
The sequence of processes in the formation of soils
is:
Weathering of rocks and minerals -> formation of
regolith or parent material ->formation of true soil from regolith
Rock ->Weathering ->Regolith ->Soil
forming factors and processes ->True soil (otherwise)
Two processes involved in the formation of soil are:
- Formation of regolith by breaking down
(weathering) of the bed rock.
- The addition of organic matter through the
decomposition of plant and animal tissues, and reorganization of these
components by soil forming processes to form soil.
Weathering: A process of disintegration and
decomposition of rocks and minerals which are brought about by physical agents
and chemical processes, leading to the formation of Regolith (unconsolidated
residues of the weathering rock on the earth’s surface or above the solid
rocks).
(OR)
The process by which the earth’s crust or
lithosphere is broken down by the activities of the atmosphere, with the aid of
the hydrosphere and biosphere
(OR)
The process of transformation of solid rocks into
parent material or Regolith
Parent material: It is the regolith or at least its upper portion.
May be defined as the unconsolidated and more or less chemically weathered
mineral materials from which soil are developed
Two basic processes of Weathering:
Physical (or) mechanical - disintegration
Chemical – decomposition
In addition, another process: Biological and all
these processes are work hand in hand
Depending up on the agents taking part in weathering processes, it is
classified into three types.
Different agents of weathering
|
Physical/ Mechanical (disintegration)
|
Chemical (decomposition)
|
Biological
(disint + decomp)
|
|
1.Physical
condition of rock
|
1.Hydration
|
1.Man &
animals
|
|
2.Change in
temperature
|
2.Hydrolysis
|
2. higher plants
& their roots
|
|
3.Action of H2O
|
3.Solution
|
3.Micro organisms
|
|
-fragment &
transport
|
4.Carbonation
|
|
|
- action of
freezing
|
5.Oxidation
|
|
|
- alter. Wet
& drying
|
6.Reduction
|
|
|
- action of
glaciers
|
|
|
|
4.Action of wind
|
|
|
|
5. Atmosp.
electric pheno
|
|
|
Physical
weathering of Rocks
The rocks are disintegrated and are broken down to
comparatively smaller pieces, with out producing any new substances
1. Physical condition of rocks: The permeability of
rocks is the most important single factor.
- Coarse textured (porous) sand stone weather
more readily than a fine textured (almost solid) basalt.
- Unconsolidated volcanic ash weather quickly as
compared to unconsolidated coarse deposits such as gravels
2. Action of Temperature: The variations in
temperature exert great influence on the disintegration of rocks.
- During day time, the rocks get heated up by
the sun and expand. At night, the temperature falls and the rocks get
cooled and contract.
- This alternate expansion and contraction
weakens the surface of the rock and crumbles it because the rocks do not
conduct heat easily.
- The minerals with in the rock also vary in
their rate of expansion and contraction
- The cubical expansion of quartz is twice as
feldspar
- Dark coloured rocks are subjected to fast
changes in temperature as compared to light coloured rocks
- The differential expansion of minerals in a
rock surface generates stress between the heated surface and cooled unexpanded
parts resulting in fragmentation of rocks.
- This process causes the surface layer to peel
off from the parent mass and the rock ultimately disintegrates. This
process is called Exfoliation
3. Action of Water: Water acts as a disintegrating,
transporting and depositing agent.
i) Fragmentation and transport: Water beats over
the surface of the rock when the rain occurs and starts flowing towards the
ocean
- Moving water has the great cutting and
carrying force.
- It forms gullies and ravines and carries with
the suspended soil material of variable sizes.
- Transporting power of water varies. It is
estimated that the transporting power of stream varies as the sixth power
of its velocity i.e the greater the speed of water, more is the
transporting power and carrying capacity
|
Speed/Sec
|
Carrying capacity
|
|
15 cm
|
fine sand
|
|
30 cm
|
gravel
|
|
1.2 m
|
stones ( 1kg)
|
|
9.0 m
|
boulders (
several tons)
|
The disintegration is greater near the source of
river than its mouth
ii) Action of freezing: Frost is much more
effective than heat in producing physical weathering
- In cold regions, the water in the cracks and
crevices freezes into ice and the volume increases to one tenth
- As the freezing starts from the top there is
no possibility of its upward expansion. Hence, the increase in volume
creates enormous out ward pressure which breaks apart the rocks
iii) Alternate wetting and Drying: Some natural
substances increase considerably in volume on wetting and shrink on
drying.(e.g.) smectite, montmorilonite
- During dry summer/ dry weather – these clays
shrink considerably forming deep cracks or wide cracks.
- On subsequent wetting, it swells.
- This alternate swelling and shrinking/ wetting
or drying of clay enriched rocks make them loose and eventually breaks
iv) Action of glaciers:
- In cold regions, when snow falls, it
accumulates and changes into ice sheet.
- These big glaciers start moving owing to the
change in temperature and/or gradient.
- On moving, these exert tremendous pressure
over the rock on which they pass and carry the loose materials
- These materials get deposited on reaching the
warmer regions, where its movement stops with the melting of ice
4. Action of wind:
- Wind has an erosive and transporting effect.
Often when the wind is laden with fine material viz., fine sand, silt or
clay particles, it has a serious abrasive effect and the sand laden winds
itch the rocks and ultimately breaks down under its force
- The dust storm may transport tons of material
from one place to another. The shifting of soil causes serious wind
erosion problem and may render cultivated land as degraded (e.g.)
Rajasthan deserts
5. Atmospheric electrical phenomenon: It is an important factor causing
break down during rainy season and lightning breaks up rocks and or widens
cracks.
Chemical
Weathering of Rocks
Decomposition of rocks and minerals by various
chemical processes is called chemical weathering. It is the most important
process for soil formation.
Chemical weathering takes place mainly at the
surface of rocks and minerals with disappearance of certain minerals and the
formation of secondary products (new materials). This is called chemical
transformation.
Feldspar + water -> clay mineral + soluble cations
and anions
Chemical weathering becomes more effective as the
surface area of the rock increases.
Since the chemical reactions occur largely on the
surface of the rocks, therefore the smaller the fragments, the greater the
surface area per unit volume available for reaction.
The effectiveness of chemical weathering is closely
related to the mineral composition of rocks. (e.g.) quartz responds far slowly
to the chemical attack than olivine or pyroxene.
Average mineralogical composition (%)
|
Composition
|
Granite
|
Basalt
|
Shale
|
S. Stone
|
L. Stone
|
|
Feldspar
|
52.4
|
46.2
|
30.0
|
11.5
|
-
|
|
Quartz
|
31.3
|
-
|
2.3
|
66.8
|
-
|
|
Pyrox - amphi
|
-
|
44.5
|
-
|
-
|
-
|
|
FeO mineral
|
2.0
|
9.3
|
10.5
|
2.0
|
-
|
|
Clay mineral
|
14.3
|
-
|
25.0
|
6.6
|
24.0
|
|
Carbonates
|
-
|
-
|
5.7
|
11.1
|
76.0
|
Chemical Processes of weathering:
1. Hydration: Chemical combination of water
molecules with a particular substance or mineral leading to a change in
structure.
Soil forming minerals in rocks do not contain any
water and they under go hydration when exposed to humid conditions. Up on
hydration there is swelling and increase in volume of minerals. The minerals
loose their luster and become soft.
It is one of the most common processes in nature and works with secondary
minerals, such as aluminium oxide and iron oxide minerals and gypsum. (e.g.)
a) 2Fe2O3 + 3HOH
-> 2Fe2O3 .3H2O
(Hematite) (Red)
(Limonite)
(Yellow)
b) Al2O3 + 3HOH
-> Al2O3 .3H2O
(Bauxite)
(Hyd. aluminium Oxide)
c) CaSO4 + 2H2O
-> CaSO4 .2H2O
(Anhydrite)
(Gypsum)
d) 3(MgO.FeO.SiO2) + 2H2O -> 3MgO.2SiO2.2H2O + SiO2 + 3H2O
(Olivine)
(Serpentine)
2. Hydrolysis: Most important process in chemical weathering. It is due to the
dissociation of H2O into H+ and OH- ions which chemically combine with
minerals and bring about changes, such as exchange, decomposition of
crystalline structure and formation of new compounds. Water acts as a weak acid
on silicate minerals.
KAlSi3O8 + H2O
->
HAlSi3O8 + KOH
(Orthoclase)
(Acid
silt clay)
HAlSi3O8 + 8 HOH -> Al2O3
.3H2O +
6 H2SiO3
(Recombination) (Hyd. Alum.
oxide) (Silicic acid)
This reaction is important because of two reasons.
- clay, bases and Silicic acid - the substances
formed in these reactions - are available to plants
- water often containing CO2 (absorbed from
atmosphere), reacts with the minerals directly to produce insoluble clay minerals,
positively charged metal ions (Ca++, Mg++, Na+, K+ ) and negatively
charged ions (OH-, HCO3-) and some soluble silica – all these ions are
made available for plant growth.
3. Solution: Some substances present in the rocks
are directly soluble in water. The soluble substances are removed by the
continuous action of water and the rock no longer remains solid and form holes,
rills or rough surface and ultimately falls into pieces or decomposes. The
action is considerably increased when the water is acidified by the dissolution
of organic and inorganic acids. (e.g) halites, NaCl
NaCl + H2O -> Na+, Cl- , H2O (dissolved ions with water)
4. Carbonation: Carbon dioxide when dissolved in
water it forms carbonic acid.
2H2O + CO2 -> H2CO3
This carbonic acid attacks many rocks and minerals and brings them into
solution. The carbonated water has an etching effect up on some rocks,
especially lime stone. The removal of cement that holds sand particles together
leads to their disintegration.
CaCO3 +
H2CO3
-> Ca(HCO3)2
(Calcite) slightly
soluble (Ca bi carbonate)
readily soluble
5. Oxidation: The process of addition and combination of oxygen to
minerals. The absorption is usually from O2 dissolved in soil water and that
present in atmosphere. The oxidation is more active in the presence of moisture
and results in hydrated oxides.(e.g) minerals containing Fe and Mg.
4FeO (Ferrous oxide) + O2 -> 2Fe2O3
(Ferric oxide)
4Fe3O4 (Magnetite) + O2
-> 6Fe2O3 (Hematite)
2Fe2O3 (Hematite) + 3H2O -> 2Fe2O3
.3H2O(Limonite)
6. Reduction: The process of removal of oxygen and is the
reverse of oxidation and is equally important in changing soil colour to grey,
blue or green as ferric iron is converted to ferrous iron compounds. Under the
conditions of excess water or water logged condition (less or no oxygen),
reduction takes place.
2Fe2O3 (Hematite) - O2 -> 4FeO( Ferrous oxide) - reduced form
In conclusion, during chemical weathering igneous and metamorphic rocks can be
regarded as involving destruction of primary minerals and the production of
secondary minerals.
In sedimentary rocks, which is made up of primary and secondary minerals,
weathering acts initially to destroy any relatively weak bonding agents (FeO) and
the particles are freed and can be individually subjected to weathering.
Biological
Weathering of Rocks
Unlike
physical and chemical weathering, the biological or living agents are
responsible for both decomposition and disintegration of rocks and minerals.
The biological life is mainly controlled largely by the prevailing environment.
1. Man and Animals:
1.
The action of man in disintegration of rocks is well known as he
cuts rocks to build dams, channels and construct roads and buildings. All these
activities result in increasing the surface area of the rocks for attack of
chemical agents and accelerate the process of rock decomposition.
2.
A large number of animals, birds, insects and worms, by their
activities they make holes in them and thus aids for weathering.
3.
In tropical and sub tropical regions, ants and termites build
galleries and passages and carry materials from lower to upper surface and
excrete acids. The oxygen and water with many dissolved substances, reach every
part of the rock through the cracks, holes and galleries, and thus brings about
speedy disintegration.
4.
Rabbits, by burrowing in to the ground, destroy soft rocks. Moles,
ants and bodies of the dead animals, provides substances which react with
minerals and aid in decaying process
5.
The earthworms pass the soil through the alimentary canal and thus
bring about physical and chemical changes in soil material.
2. Higher Plants and Roots: The roots of
trees and other plants penetrates into the joints and crevices of the rocks. As
they grew, they exert a great disruptive force and the hard rock may break
apart. (e.g.) pipal tree growing on walls/ rocks
The
grass root form a sponge like mass prevents erosion and conserve moisture and
thus allowing moisture and air to enter in to the rock for further action.
Some
roots penetrate deep into the soil and may open some sort of drainage channel.
The roots running in crevices in lime stone and marble produces acids. These
acids have a solvent action on carbonates.
The
dead roots and plant residues decompose and produce carbon dioxide which is of
great importance in weathering.
3.
Micro- organisms: In early stages of mineral decomposition and
soil formation, the lower forms of plants and animals like, mosses, bacteria
and fungi and actinomycetes play an important role. They extract nutrients from
the rock and N from air and live with a small quantity of water. In due course
of time, the soil develops under the cluster of these micro-organisms.
This
organism closely associated with the decay of plant and animal remains and thus
liberates nutrients for the use of next generation plants and also produces CO2
and organic compounds which aid in mineral decomposition.
Weathering of
Minerals
There
are many factors which influence the weathering of minerals.
1.
1. Climatic conditions
2.
2. Physical characteristics
3.
3. Chemical and structural characteristics
1. Climatic Conditions: The climatic
condition, more than any other factor tends to control the kind and rate of
weathering. Under conditions of low rainfall, there is a dominance of physical
weathering which reduces the size and increases the surface area with little
change in volume.
The
increase in moisture content encourages chemical as well as mechanical changes
and new minerals and soluble products are formed.
The
weathering rates are generally fastest in humid tropical regions as there is
sufficient moisture and warmth to encourage chemical decomposition.
The
easily weather able minerals disappear on account of intense chemical
weathering and more resistant products (hydrous oxides of Fe and Al) tend to
accumulate
Climate
controls the dominant type of vegetation which in turn controls the bio
chemical reactions in soils and mineral weathering.
2. Physical Characteristics:
i) Differential composition
ii) Particle size
iii) Hardness and degree of cementation
3. Chemical and structural characteristics:
Chemical: For minerals of given particle size, chemical and crystalline
characteristics determine the ease of decomposition. (e.g.) gypsum – sparingly
soluble in water, is dissolved and removed in solution form under high
rainfall.
Ferro magnesium minerals are more susceptible to chemical weathering than
feldspar and quartz
Tightness of packing of ions in crystals: Less tightly packed minerals like olivine
and biotitic are easily weathered as compared to tightly packed zircon and
muscovite (resistant)
Chemical
weathering of Silicates
The most important silicates are quartz, feldspar
and certain Ferro magnesium minerals
|
Weathering products of common silicate minerals
|
|
Minerals
|
Composition
|
Decomposition products
|
|
Minerals
|
Others
|
|
Olivine
|
(Fe, Mg)2SiO2
|
serpentine,
limonitehaematite, quartz
|
Some Si in
solution, carbonates of Fe and Mg
|
|
Pyroxenes
|
Fe, Mg
|
Clay, calcite,
limonite
|
Some Si in
solution, carbonates of Ca and Mg
|
|
Amphibole
|
Ca- silicates
|
Haematite, quartz
|
-do -
|
|
Biotite
|
Al
|
|
-do -
|
|
Plagioclase
|
Calcic
|
Clay, quartz,
calcite
|
Some Si in
solution, Na and Ca carbonates
|
|
Sodic
|
|
-do-
|
|
Orthoclase
|
Potassic
|
Clay, quartz
|
Some Si in
solution, potassium carbonate
|
|
Quartz
|
|
Quartz grains
|
Some Si in
solution
|
Soil Forming
Factors
The
soil formation is the process of two consecutive stages.
1.
The weathering of rock (R) into Regolith
2.
The formation of true soil from Regolith
The
evolution of true soil from regolith takes place by the combined action of soil
forming factors and processes.
1.
The first step is accomplished by weathering (disintegration &
decomposition)
2.
The second step is associated with the action of Soil Forming
Factors
Weathering Factors
Dokuchaiev (1889) established that the soils develop as a result of the action
of soil forming factors
S = f ( P, Cl, O )
Further,
Jenny (1941) formulated the following equation
S = f (Cl, O, R, P, T, …)
Where,
Cl – environmental climate
o – Organisms and vegetation (biosphere)
r – Relief or topography
p – Parent material
t- Time
… - additional unspecified factors
The five soil forming factors, acting simultaneously at any point on the
surface of the earth, to produce soil
Two groups – Passive i) Parent material, ii) Relief, iii) Time
Active IV) Climate, v) Vegetation & organism
Passive
Soil forming factors
The passive soil forming factors are those which
represent the source of soil forming mass and conditions affecting it. These
provide a base on which the active soil forming factors work or act for the
development of soil.
Parent Material: It is that mass (consolidated material) from which the
soil has formed.
Two groups of parent material
i) Sedentary: Formed in original place. It is the residual parent material. The
parent material differ as widely as the rocks
ii) Transported: The parent material transported from their place of origin.
They are named according to the main force responsible for the transport and
redeposition.
a) By gravity - Colluvial
b) By water - Alluvial, Marine, Locustrine
c) By ice - Glacial
d) By wind – Eolian
Colluvium: It is the poorly sorted materials near the base of strong
slopes transported by the action of gravity.
Alluvium: The material transported and deposited by water is, found
along major stream courses at the bottom of slopes of mountains and along small
streams flowing out of drainage basins.
Locustrine: Consists of materials that have settled out of the quiet
water of lakes.
Moraine: Consists of all the materials picked up, mixed,
disintegrated, transported and deposited through the action of glacial ice or
of water resulting primarily from melting of glaciers.
Loess or Aeolian: These are the wind blown materials.
When the texture is silty - loess; when it is sand - Eolian.
The soils developed on such transported parent materials bear the name of the
parent material; viz. Alluvial soils from alluvium, Colluvial soils from
Colluvium etc. In the initial stages, however, the soil properties are mainly
determined by the kind of parent material.
Endodynamomorphic
soils: With advanced
development and excessive leaching, the influence of parent material on soil
characteristics gradually diminishes. There are soils wherein the composition
of parent material subdues the effects of climate and vegetation. These soils
are temporary and persist only until the chemical decomposition becomes active
under the influence of climate and vegetation.
Ectodynamomorphic soils: Development of normal profile under the
influence of climate and vegetation.
Soil properties as influenced by parent material: Different parent materials
affect profile development and produce different soils, especially in the
initial stages.
- Acid igneous rocks (like granite, rhyolite)
produce light-textured soils (Alfisols).
- Basic igneous rocks (basalt), alluvium or
Colluvium derived from limestone or basalt, produce fine-textured
cracking-clay soils (Vertisols).
- Basic alluvium or Aeolian materials produce
fine to coarse-textured soils (Entisols or Inceptisols).
- The nature of the elements released during the
decaying of rocks has a specific role in soil formation. (e.g.) Si and Al
form the skeleton for the production of secondary clay minerals.
- Iron and manganese are important for imparting
red colour to soils and for oxidation and reduction phenomena.
- Sodium and potassium are important dispersing
agents for day and humus colloids.
- Calcium and magnesium have a flocculating
effect and result in favorable and stable soil structure for plant growth.
2. Relief or
Topography: The relief and
topography sometimes are used as synonymous terms. They denote the
configuration of the land surface. The topography refers to the differences in
elevation of the land surface on a broad scale.
The prominent types of topography designations, as given in FAO Guidelines
(1990) are:
|
Sr. No
|
Land surface
|
with slopes of
|
|
1
|
Flat to Almost
flat
|
0 – 2 %
|
|
2
|
Gently undulating
|
2 - 5 %
|
|
3
|
Undulating
|
5 – 10 %
|
|
4
|
Rolling
|
10 – 15 %
|
|
5
|
Hilly
|
15 –3 0 %
|
|
6
|
Steeply dissect
|
>30% with
moderate range of
elevation( <300 m)
|
|
7
|
Mountainous
|
> 30% with
great range of
elevation(>300 m)
|
Soil formation on
flat to almost flat position: On level
topographic positions, almost the entire water received through rain percolates
through the soil. Under such conditions, the soils formed may be considered as
representative of the regional climate. They have normal solum with distinct
horizons. But vast and monotonous level land with little gradient often has
impaired drainage conditions.
Soil formation on undulating topography: The soils on steep slopes are
generally shallow, stony and have weakly- developed profiles with less distinct
horizonation. It is due to accelerated erosion, which removes surface material
before it has the time to develop. Reduced percolation of water through soil is
because of surface runoff, and lack of water for the growth of plants, which
are responsible for checking of erosion and promote soil formation.
Soil formation in depression: The depression areas in semi-arid and sub
humid regions reflect more moist conditions than actually observed on level
topographic positions due to the additional water received as runoff. Such
conditions (as in the Tarai region of the Uttar Pradesh) favour more vegetative
growth and slower rate of decay of organic remains. This results in the
formation of comparatively dark- coloured soils rich in organic matter
(Mollisols).
Soil formation and Exposure/ Aspect: Topography affects soil formation
by affecting temperature and vegetative growth through slope exposures
(aspect}. The southern exposures (facing the sun) are warmer and subject to
marked fluctuations in temperature and moisture. The northern exposures, on the
other hand are cooler and more humid. The eastern and western exposures occupy
intermediate position in this respect.
3. Time: Soil formation is a very slow process requiring thousands of
years to develop a mature pedon. The period taken by a given soil from the
stage of weathered rock (i.e. regolith) up to the stage of maturity is
considered as time. The matured soils mean the soils with fully developed
horizons (A, B, C). It takes hundreds of years to develop an inch of soil. The
time that nature devotes to the formation of soils is termed as Pedological
Time.
It has been observed that rocks and minerals
disintegrate and/or decompose at different rates; the coarse particles of
limestone are more resistant to disintegration than those of sandstone.
However, in general, limestone decomposes more readily than sandstone (by
chemical weathering).
|
Weathering stages in soil formation
|
|
Sr. No
|
Stages
|
Characteristic
|
|
1
|
Initial
|
Un weathered
parent material
|
|
2
|
Juvenile
|
Weathering
started but much of the original
material still un weathered
|
|
3
|
Virile
|
Easily weather
able minerals fairly decomposed; clay content increased, slowly weather able
minerals still appreciable
|
|
4
|
Senile
|
Decomposition
reaches at a final stage; only
most resistant minerals survive
|
|
5
|
Final
|
Soil development
completed under prevailing
environments
|
- The soil properties also change with time, for
instance nitrogen and organic matter contents increase with time provided
the soil temperature is not high.
- CaCO3 content may decrease or even lost with
time provided the climatic conditions are not arid
- In humid regions, the H+ concentration
increases with time because of chemical weathering.
Active Soil
Forming Factors
The active soil forming factors are those which
supply energy that acts on the mass for the purpose of soil formation. These
factors are climate and vegetation (biosphere).
1. Climate: Climate is the most significant factor controlling the type and
rate of soil formation. The dominant climates recognized are:
- Arid climate: The precipitation here is far
less than the water-need. Hence the soils remain dry for most of the time
in a year.
- Humid climate: The precipitation here is much
more than the water need. The excess water results in leaching of salt and
bases followed by translocation of clay colloids.
- Oceanic climate: Moderate seasonal variation
of rainfall and temperature.
- Mediterranean climate: The moderate
precipitation. Winters and summers are dry and hot.
- Continental climate: Warm summers and extremely
cool or cold winters.
- Temperate climate: Cold humid conditions with
warm summers.
- Tropical and subtropical climate: Warm to hot
humid with isothermal conditions in the tropical zone.
Climate affects the soil formation directly and
indirectly.
Directly, climate affects the soil formation by supplying water and heat to
react with parent material.
Indirectly, it determines the fauna and flora
activities which furnish a source of energy in the form of organic matter. This
energy acts on the rocks and minerals in the form of acids, and salts are
released. The indirect effects of climate on soil formation are most clearly
seen in the relationship of soils to vegetation.
Precipitation and temperature are the two major
climatic elements which contribute most to soil formation.
Precipitation: Precipitation is the most important among the
climatic factors. As it percolates and moves from one part of the parent
material to another. It carries with it substances in solution as well as in
suspension. The substances so carried are re deposited in another part or
completely removed from the material through percolation when the soil moisture
at the surface evaporates causing an upward movement of water. The soluble
substances move with it and are translocated to the upper layer. Thus rainfall
brings about a redistribution of substances both soluble as well as in
suspension in soil body.
Temperature:
- Temperature is another climatic agent
influencing the process of soil formation.
- High temperature hinders the process of
leaching and causes an upward movement of soluble salts.
- High temperature favors rapid decomposition of
organic matter and increase microbial activities in soil while low
temperatures induce leaching by reducing evaporation and there by favour
the accumulation of organic matter by slowing down the process of
decomposition. Temperature thus controls the rate of chemical and
biological reactions taking place in the parent material.
Jenney (1941) computed that in the tropical regions the rate of
weathering proceeds three times faster than in temperate regions and nine times
faster than in arctic.
2. Organism & Vegetation
Organism:
- The active components of soil ecosystem are
plants, animals, microorganisms and man.
- The role of microorganisms in soil formation
is related to the humification and mineralization of vegetation
- The action of animals especially burrowing
animals to dig and mix-up the soil mass and thus disturb the parent
material
- Man influences the soil formation through his
manipulation of natural vegetation, agricultural practices etc.
- Compaction by traffic of man and animals
decrease the rate of water infiltration into the soil and thereby increase
the rate of runoff and erosion.
Vegetation:
- The roots of the plants penetrate into the
parent material and act both mechanically and chemically.
- They facilitate percolation and drainage and
bring about greater dissolution of minerals through the action of CO2 and
acidic substances secreted by them.
- The decomposition and humification of the
materials further adds to the solubilization of minerals
- Forests – reduces temperature, increases
humidity, reduce evaporation and increases precipitation.
- Grasses reduce runoff and result greater
penetration of water in to the parent material.
Soil Forming
Processes
The
pedogenic processes, although slow in terms of human life, yet work faster than
the geological processes in changing lifeless parent material into true soil
full of life.
·
The pedogenic processes are extremely complex and dynamic
involving many chemical and biological reactions, and usually operate
simultaneously in a given area.
·
One process may counteract another, or two different processes may
work simultaneously to achieve the same result.
·
Different processes or combination of processes operate under
varying natural environment.
The
collective interaction of various soil forming factors under different
environmental conditions set a course to certain recognized soil forming
processes.
The basic process involved in soil formation (Simonson, 1959) includes the following.
·
Gains or Additions of water, mostly as rainfall, organic and
mineral matter to the soil.
·
Losses of the above materials from the soil.
·
Transformation of mineral and organic substances within the soil.
·
Translocation or the movement of soil materials from one point to
another within the soil. It is usually divided into
1.
movement of solution (leaching) and
2.
movement in suspension (eluviation) of clay, organic matter and
hydrous oxides
3.
Fundamental Soil forming Processes
4.
Humification: Humification is the process of transformation of raw
organic matter into humus. It is extremely a complex process involving various
organisms.
First, simple compounds such as sugars and starches are attacked followed by
proteins and cellulose and finally very resistant compounds, such as tannins,
are decomposed and the dark coloured substance, known as humus, is formed.
Eluviation: It is the mobilization and translocation of certain constituent’s
viz. Clay, Fe2O3, Al2O3, SiO2, humus, CaCO3, other salts etc. from one point of
soil body to another. Eluviation means washing out. It is the process of
removal of constituents in suspension or solution by the percolating water from
the upper to lower layers. The eluviation encompasses mobilization and
translocation of mobile constituents resulting in textural differences. The
horizon formed by the process of eluviation is termed as eluvial horizon (A2 or
E horizon).
Translocation depends upon relative mobility of elements and depth of
percolation.
Illuviation: The process of deposition of soil materials (removed from the
eluvial horizon) in the lower layer (or horizon of gains having the property of
stabilizing translocated clay materials) is termed as Illuviation. The horizons
formed by this process are termed as illuvial horizons (B-horizons, especially
Bt) The process leads to textural contrast between E and Bt horizons, and
higher fine: total clay ratio in the Bt horizon.
Horizonation: It is the process of differentiation of soil in different
horizons along the depth of the soil body. The differentiation is due to the
fundamental processes, humification, eluviation and Illuviation.
Specific
Soil Forming Processes
The basic pedologic processes provide a framework
for later operation of more specific processes
Calcification: It is the process of precipitation and accumulation of calcium
carbonate (CaCO3) in some part of the profile. The accumulation of CaCO3 may
result in the development of a calcic horizon. Calcium is readily soluble
in acid soil water and/or when CO2 concentration is high in root zone as:
CO2 + H2O --> H2CO3
H2CO3 + Ca --> Ca (HCO3)2 (soluble)
Temp.
Ca (HCO3)2 --> CaCO3 + H2O + CO2 (precipitates)
CO2
The process of precipitation after mobilization under these conditions is
called calcification and the resulting illuviated horizon of carbonates is
designated as Bk horizon (Bca).
Decalcification: It is the reverse of calcification that is the process
of removal of CaCO3 or calcium ions from the soil by leaching
Temp.
CaCO3 + CO2 + H2O ---> Ca (HCO3)2 (soluble)
(insoluble)
CO2
Podzolization: It is a process of soil formation
resulting in the formation of Podzols and Podzolic soils. In many respects,
podzolization is the negative of calcification. The calcification process tends
to concentrate calcium in the lower part of the B horizon, whereas
podzolization leaches the entire solum of calcium carbonates.
Apart from calcium, the other bases are also removed
and the whole soil becomes distinctly acidic. In fact, the process is
essentially one of acid leaching. The process operates under favorable
combination of the following environments.
i) Climate: A cold and humid climate is most favorable for podzolization
ii) Parent material: Siliceous (Sandy) material, having poor reserves of
weather able minerals, favor the operation of podzolization as it helps in easy
percolation of water.
iii) Vegetation: Acid producing vegetation such as coniferous pines is essential
iv) Leaching and Translocation of Sesquioxide: In the process of decomposition
of organic matter various organic acids are produced. The organic acids thus
formed act with Sesquioxide and the remaining clay minerals, forming organic-
Sesquioxide and organic clay complexes, which are soluble and move with the
percolating water to the lower horizons (Bh, Bs). Aluminium ions in a water
solution hydrolyze and make the soil solution very acidic.
2Al +6H2O à2 Al(OH)3 + 6H+
As iron and aluminium move about, the A horizon
gives a bleached grey or ashy appearance. The Russians used the term Podzols
(pod means under, the Zola means ash like i.e. ash-like horizon appearing
beneath the surface horizon) for such soils.
To conclude, the Podzolization is a soil forming
process which prevails in a cold and humid climate where coniferous and acid
forming vegetations dominate. The humus and Sesquioxide become mobile and
leached out from the upper horizons and deposited in the lower horizon.
4. Laterization: The term laterite is derived from
the word later meaning brick or tile and was originally applied to a group of
high clay Indian soils found in Malabar hills of Kerala, Tamil Nadu, Karnataka
and Maharashtra.
It refers specifically to a particular cemented
horizon in certain soils which when dried, become very hard, like a brick. Such
soils (in tropics) when massively impregnated with Sesquioxide (iron and
aluminium oxides) to extent of 70 to 80 per cent of the total mass, are called
laterite or latosols (Oxisols). The soil forming process is called Laterization
or Latozation.
Laterization is the process that removes silica,
instead of sesquioxides from the upper layers and thereby leaving sesquioxides
to concentrate in the solum. The process operates under the following
conditions.
i) Climate: Unlike podzolization, the process of
laterization operates most favorable in warm and humid (tropical) climate with
2000 to 2500 mm rainfall and continuous high temperature (25°C) throughout the
year.
ii) Natural vegetation: The rain forests of
tropical areas are favorable for the process.
iii) Parent Material: Basic parent materials,
having sufficient iron bearing ferromagnesian minerals (Pyroxene, amphiboles,
biotite and chlorite), which on weathering release iron, are congenial for the
development of laterites.
5. Gleization: The term glei is of Russian origin means blue,
grey or green clay. The Gleization is a process of soil formation resulting in
the development of a glei (or gley horizon) in the lower part of the soil
profile above the parent material due to poor drainage condition (lack of
oxygen) and where waterlogged conditions prevail. Such soils are called hydro
orphic soils.
The process is not particularly dependent on
climate (high rainfall as in humid regions) but often on drainage conditions.
The poor drainage conditions result from:
- Lower topographic position, such as depression
land, where water stands continuously at or close to the surface.
- Impervious soil parent material, and.
- Lack of aeration.
Under such conditions, iron compounds are reduced
to soluble ferrous forms. The reduction of iron is primarily biological and
requires both organic matter and microorganisms capable of respiring
anaerobically. The solubility of Ca, Mg, Fe, and Mn is increased and most of
the iron exists as Fe++ organo complexes in solution or as mixed precipitate of
ferric and ferrous hydroxides.
This is responsible for the production of typical
bluish to grayish horizon with mottling of yellow and or reddish brown colors.
6. Salinization: It is the process of accumulation of salts, such as
sulphates and chlorides of calcium, magnesium, sodium and potassium, in soils
in the form of a salty (salic) horizon. It is quite common in arid and semi
arid regions. It may also take place through capillary rise of saline ground
water and by inundation with seawater in marine and coastal soils. Salt
accumulation may also result from irrigation or seepage in areas of impeded
drainage.
7. Desalinization: It is the removal by leaching of excess soluble
salts from horizons or soil profile (that contained enough soluble salts to
impair the plant growth) by ponding water and improving the drainage conditions
by installing artificial drainage network.
8. Solonization or
Alkalization: The process
involves the accumulation of sodium ions on the exchange complex of the clay,
resulting in the formation of sodic soils (Solonetz). All cations in solution
are engaged in a reversible reaction with the exchange sites on the clay and
organic matter particles.
The reaction can be represented as:
Ca.Mg.2NaX à Ca++ +Mg++ +2Na+ + x-6 +3CO3 2- à Na2CO3 + MgCO3
+CaCO3
(Where X represents clay or organic matter exchange sites)
9. Solodization or
dealkalization: The process refers
to the removal of Na+ from the exchange sites. This process involves dispersion
of clay. Dispersion occurs when Na+ ions become hydrated. Much of the
dispersion can be eliminated if Ca++ and or Mg++ ions are concentrated in the
water, which is used to leach the soonest. These Ca and Mg ion can replace the
Na on exchange complex, and the salts of sodium are leached out as:
2NaX + CaSO4 à Na2SO4 + CaX
(leachable)
10. Pedoturbation: Another process that may be operative in soils is
pedoturbation. It is the process of mixing of the soil. Mixing to a certain
extent takes place in all soils. The most common types of pedoturbation are:
- Faunal pedoturbation: It is the mixing of soil
by animals such as ants, earthworms, moles, rodents, and man himself
- Floral pedoturbation : It is the mixing of
soil by plants as in tree tipping that forms pits and mounds
- Argillic pedoturbation: It is the mixing of
materials in the solum by the churning process caused by swell shrink
clays as observed in deep Black Cotton Soils.
Development
of Soil Profile
The development of soil profile is a constructive
process where in disintegrated material resulted from weathering of rocks and
minerals gets converted into a soil body.
Definition of soil profile: The vertical section of
the soil showing the various layers from the surface to the unaffected parent
material is known as a soil profile.
The various layers are known as horizons. A soil
profile contains three main horizons.
They are named as horizon A, horizon B and horizon
C.
- The surface soil or that layer of soil at the
top which is liable to leaching and from which some soil constituents have
been removed is known as horizon A or the horizon of eluviation.
- The intermediate layer in which the materials
leached from horizon A have been re-deposited is known as horizon B or the
horizon of illuviation.
- The parent material from which the soil is
formed is known as horizon C.
A Study of soil profile is important as it is
historic record of all the soil forming processes and it forms the basis for
the study in pedagogical investigations.
Soil profile is the key for the soil classification and also forms the basis
for the practical utility of soils.
A hypothetical mineral soil profile will include O, A, B, C and R master
horizons and all the possible sub-horizons.
Master horizons and sub horizons
O horizon: It is called as organic horizon. It is
formed in the upper part of the mineral soil, dominated by fresh or partly
decomposed organic materials.
- This horizon contains more than 30% organic
matter if mineral fraction has more than 50 % clay (or) more than 20 %
organic matter if mineral fraction has less clay.
- The organic horizons are commonly seen in
forest areas and generally absent in grassland, cultivated soils.
- O1 - Organic horizon in which
the original forms of the plant and animal residues can be recognized
through naked eye.
- O2 - Organic horizon in which
the original plant or animal matter can not be recognized through naked
eye.
- A horizon - Horizon of organic matter
accumulation adjacent to surface and that has lost clay, iron and
Aluminium.
- A1 - Top most mineral horizon
formed adjacent to the surface. There will be accumulation of humified
organic matter associated with mineral fraction and darker in Colour
than that of lower horizons due to organic matter.
- A2 - Horizon of maximum
eluviation of clay, iron and aluminium oxides and organic matter. Loss
of these constituents generally results in accumulation of quartz and
other sand and silt size resistant minerals. Generally lighter in
Colour than horizons above and below.
- A3 - A transitional layer
between A and B horizons with more dominated properties of A1 or A2
above than the underlying B horizon. This horizon is sometimes absent.
Solum.
- B horizon - Horizon in which the dominant
features are accumulation of clay, iron, aluminium or humus alone or in
combination. Coating of sesquioxides will impart darker, stronger of red
Colour than overlying or underlying horizons.
- B1 - A transitional layer
between A and B. More like A than B.
- B2 - Zone of maximum
accumulation of clay, iron and aluminium oxide that may have moved down
from upper horizons or may have formed in situ. The organic matter
content is generally higher and Colour darker than that of A2 horizon
above.
- B3 - Transitional horizon
between B and C and with properties more similar to that of overlying
B2 than underlying C.
- C horizon - It is the horizon below the solum
(A + B), relatively less affected by soil forming processes. It is outside
the zone of major biological activity. It may contain accumulation of
carbonates or sulphates, calcium and magnesium
- R - Underlying consolidated bed rock and it
may or may not be like the parent rock from which the solum is formed.
Besides, lower case letters are used to indicate
the special features of master horizons. This case letters follow the
subdivisions of master horizons. E.g. Ap - ploughed layer, B2t - illuvial clay
When two or more genetically unrelated (contrasting) materials are present in a
profile as in the case of alluvial or colluvial soils then the phenomenon is
known as lithological discontinuity. This is indicated by the use of Roman
letters as prefixes to the master horizons. E.g. Ap, B2, II B22, IIIC.
Special Features:
Soil Individual or Polypedon: The Soil Survey Staff
(1960) defined the soil individual or polypedon (Pedon, Ground) as a natural
unit of soil that differs from its adjoining unit on the landscape in one or
more properties.
The term pedon has been proposed for small basic soil entities that are part of
the continuum mantling the land.
A pedon is the smallest volume that can be called
"a soil". The set of pedon must fit within the range of one series
and occur in a contiguous group to form a polypedon.
A polypedon is therefore, defined as a contiguous
similar pedons bounded on all sides by "not-soil or by pedons of unlike
characters. It is a real physical soils body which has a minimum area of more
than 1 sq. km and an unspecified maximum area.
Physical
Properties of Soil
Physical properties (mechanical behaviour) of a
soil greatly influence its use and behaviour towards plant growth. The plant
support, root penetration, drainage, aeration, retention of moisture, and plant
nutrients are linked with the physical condition of the soil. Physical
properties also influence the chemical and biological behaviour of soil. The
physical properties of a soil depend on the amount, size, shape, arrangement
and mineral composition of its particles. These properties also depend on
organic matter content and pore spaces
Important physical properties of soils
1. Soil texture
2. Soil structure
3. Surface area
4. Soil density,
5. Soil porosity
6. Soil colour
7. Soil consistence
Soil Texture
Definition of Soil Texture: Soil texture refers to
the relative proportion of particles or it is the relative percentage by weight
of the three soil separates viz., sand, silt and clay or simply refers to the
size of soil particles.
The proportion of each size group in a given soil (the texture) can not be
easily altered and it is considered as a basic property of a soil.
The soil separates are defined in terms of diameter in millimeters of the
particles. Soil particles less than 2 mm in diameter are excluded from soil
textural determinations.
Stones and gravels may influence the use and management of land because of
tillage difficulties but these larger particles make little or no contribution
to soil properties such as WHC and capacity to store plant nutrients and their
supply.
Gravels: 2 - 4 mm
Pebbles: 4 - 64 mm
Cobbles: 64 - 256 mm
Boulders: > 256 mm
Particles less than 2 mm are called fine earth,
normally considered in chemical and mechanical analysis.
The components of fine earth: Sand, Silt and Clay
(Soil separates. The size limits of these fractions have been established by
various organizations. There are a number of systems of naming soil separates.
- (a) The American system developed by USDA
- (b) The English system or British system ( BSI
)
- (c) The International system (ISSS)
- (d) European system
i) USDA
|
Soil
separates
|
Diameter
in mm
|
|
Clay
|
< 0.002
|
|
Silt
|
0.002 - 0.05
|
|
Very Fine Sand
|
0.05 - 0.10
|
|
Fine Sand
|
0.10 - 0.25
|
|
Medium Sand
|
0.25 - 0.50
|
|
Coarse Sand
|
0.50 - 1
|
|
Very Coarse Sand
|
1 - 2 mm
|
ii) BSI
|
Soil
separates
|
Diameter
in mm
|
|
Clay
|
< 0.002
|
|
Fine Silt
|
0.002 - 0.01
|
|
Medium Silt
|
0.01 - 0.04
|
|
Coarse Silt
|
0.04 - 0.06
|
|
Fine Sand
|
0.06 - 0.20
|
|
Medium Sand
|
0.20 - 1
|
|
Coarse Sand
|
1 - 2 mm
|
iii) ISSS
|
Soil
separates
|
Diameter
in mm
|
|
Clay
|
< 0.002
|
|
Silt
|
0.002 - 0.02
|
|
Fine Sand
|
0.02 - 0.2
|
|
Coarse Sand
|
0.2 - 2
|
iv) European System
|
Soil separates
|
Diameter
(mm)
|
|
Fine clay
|
< 0.0002 mm
|
|
Medium clay
|
0.0002 – 0.0006
|
|
Coarse clay
|
0.0006 – 0.002
|
|
Fine silt
|
0.002 - 0.006
|
|
Medium silt
|
0.006 - 0.02
|
|
Coarse silt
|
0.02 - 0.06
|
|
Fine sand
|
0.06 - 0.20
|
|
Medium sand
|
0.20 - 0.60
|
|
Coarse sand
|
0. 60 - 2.00
|
Sand:
- Usually consists of quartz but may also
contain fragments of feldspar, mica and occasionally heavy minerals viz.,
zircon, Tourmaline and hornblende.
- Has uniform dimensions
- Can be represented as spherical
- Not necessarily smooth and has jagged surface
Silt:
- Particle size intermediate between sand and
clay
- Since the size is smaller, the surface area is
more
- Coated with clay
- Has the physico- chemical properties as that
of clay to a limited extent
- Sand and Silt forms the SKELETON
Clay:
- Particle size less than 0.002 mm
- Plate like or needle like in shape
- Belong to alumino silicate group of minerals
- Some times considerable concentration of fine
particles which does not belong to alumino silicates. (e.g.) iron oxide
and CaCO3
- These are secondary minerals derived from
primary minerals in the rock
- Flesh of the soil
Knowledge on Texture is important. It is a guide to
the value of the land. Land use capability and methods of soil management
depends on texture.
Particle size distribution/ determination
The determination of relative distribution of the ultimate or individual soil
particles below 2 mm diameter is called as Particle size analysis or Mechanical
analysis
Two steps are involved
i) Separation of all the particles from each other ie. Complete dispersion into
ultimate particles
ii) Measuring the amount of each group
Separation
|
Sr. No.
|
Aggregating
agents
|
Dispersion
method
|
|
1
|
Lime and Oxides
of Fe & Al
|
Dissolving in HCl
|
|
2
|
Organic matter
|
Oxidises with
H2O2
|
|
3
|
High conc. of
electrolytes
( soluble salts)
|
Precipitate and
decant or filter with suction
|
|
4
|
Surface tension
|
Elimination of
air by stirring with water or boiling
|
After removing the cementing agents, disperse by
adding NaOH
Measurement
Once the soil particles are dispersed into ultimate particles, measurement can
be done
i) Coarser fractions - sieving - sieves used in the mechanical analysis
corresponds to the desired particle size separation for 2 mm, 1 mm and 0.5 mm –
sieves with circular holes, for smaller sizes, wire mesh screens are used
(screening)
ii) Finer fractions - by settling in a medium the
settling or the velocity of the fall of particles is influenced by viscosity of
the medium. Difference in density between the medium and falling particles,
size and shape of object
Stokes' Law:
Particle size analysis is based on a simple
principle i.e. "when soil particles are suspended in water they tend to
sink. Because there is little variation in the density of most soil particles,
their velocity (V) of settling is proportional to the square of the radius 'r'
of each particle.
Thus V = kr2, where k is a constant. This equation is referred to as Stokes'
law.
Stokes (1851) was the first to suggest the relationship between the radius of
the particles and its rate of fall in a liquid. He stated that "the
velocity of a falling particle is proportional to the square of the radius and
not to its surface. The relation between the diameter of a particle and its
settling velocity is governed by Stokes' Law:
V = 2/9 gr^2 (ds – dw) / n
Where,
V - Velocity of settling particle (cm/sec.)
g - Acceleration due to gravity cm/ sec2 (981)
ds - Density of soil particle (2.65)
dw - Density of water (1)
n - Coefficient of viscosity of water (0.0015 at 4oC)
r - Radius of spherical particles (cm).
Assumptions and Limitations of Stokes' Law
Particles are rigid and spherical / smooth. This
requirement is very difficult to fulfill, because the particles are not
completely smooth over the surface and spherical. It is established that the
particles are not spherical and irregularly shaped such as plate and other
shapes.
The particles are large in comparison with the molecules of the liquid so that
in comparison with the particle the medium can be considered as homogenous i.e.
the particles must be big enough to avoid Brownian movement. The particles less
than 0.0002 mm exhibit this movement so that the rate of falling is varied.
The fall of the particles is not hindered or affected by the proximity (very
near) of the wall of the vessel or of the adjacent particles. Many fast falling
particles may drag finer particles down along with them.
The density of the particles and water and as well
as the viscosity of the medium remain constant. But this is usually not so
because of their different chemical and mineralogical composition.
The suspension must be still. Any movement in the suspension will alter the
velocity of fall and such movement is brought by the sedimentation of larger
particles (> 0.08 mm). They settle so fast and create turbulence in the
medium.
The temperature should be kept constant so that convection currents are not set
up.
Methods of Textural determination
Numerous methods for lab and field use have been
developed
i) Elutriation method – Water & Air
ii)Pipette method
iii) Decantation/ beaker method
iv) Test tube shaking method
v) Feel method - Applicable to the field - quick method - by feeling the soil
between thumb and fingers
Feel Method
Evaluated by attempting to squeeze the moistened
soil into a thin ribbon as it is pressed with rolling motion between thumb and
pre finger or alternately to roll the soil into a thin wire
η four aspects to be seen - i) Feel by fingers, ii) Ball formation, iii)
Stickiness and iv) Ribbon formation
Soil Textural
Classes
To
convey an idea of the textural make up of soils and to give an indication of
their physical properties, soil textural class names are used. These are
grouped into three main fractions viz., Sand, Silt and Clay.
According to the proportion of these three fractions a soil is given a name to
indicate its textural composition. Such a name gives an idea not only of the
textural composition of a soil but also of its various properties in general.
On
this basis soils are classified into various textural classes like sands clays,
silts, loams etc
Sands: The sand group includes all soils in which the
sand separates make up at least 70% and the clay separate 15% or less of the
material by weight. The properties of such soils are therefore
characteristically those of sand in contrast to the stickier nature of clays.
Two specific textural classes are recognized in this group sandy and loamy sand
although in practice two subclasses are also used Loamy fine sand and loamy
very fine sand.
Silt: The silt group includes soils with at least 80%
silt and 12% or less clay. Naturally the properties of this group are dominated
by those of silt. Only one textural class - Silt is included in this group.
Clays: To be designated clay a soi1 must contain at least 35% of the clay
separate and in most cases not less than 40%. In such soils the characteristics
of the clay separates are distinctly dominant, and the class names are clay,
sandy clay and silty clay. Sandy clays may contain more sand than clay.
Likewise, the silt content of silty clays usually exceeds clay fraction.
Loams: The loam group, which contains many
subdivisions, is a more complicated soil textural class. An ideal loam may be
defined as a mixture of sand, silt and day particles that exhibits the
properties of those separates in about equal proportions. Loam soils do not
exhibit dominant physical properties of sand, silt or clay. Loam does not
contain equal percentage of sand, silt and clay. However, exhibit approximately
equal properties of sand, silt and clay.
Determination of Textural Class: In the
American system as developed by the United State Department of Agriculture
twelve textural classes are proposed.
The textural triangle: It is used
to determine the soil textural name after the percentages of sand, silt, and
clay are determined from a laboratory analysis. Since the soil's textural
classification includes only mineral particles and those of less than 2mm
diameter, the sand plus silt plus clay percentages equal 100 percent. (Note
that organic matter is not included.) Knowing the amount of any two fractions
automatically fixes the percentage of the third one.
To use the diagram, locate the percentage of clay first and project inward
parallel to sand line. Do likewise for the per cent silt and project inward
parallel to clay line and for sand, project inward parallel to silt. The point
at which the projections cross or intersect will identify the class name.
Some
times, the intersecting point exactly falls on the line between the textural
classes. Then it is customary to use the name of the finer fraction when it
happens. (e.g). Soil containing 40% clay, 30% sand and 30% silt - called as
clay rather than clay loam.
Importance of Soil Texture
Presence
of each type of soil particles makes its contribution to the nature and
properties of soil as a whole
·
Texture has good effect on management and productivity of soil.
Sandy soils are of open character usually loose and friable.
·
Such type of the texture is easy to handle in tillage operations.
·
Sand facilitates drainage and aeration. It allows rapid
evaporation and percolation.
·
Sandy soils have very little water holding capacity. Such soils
can not stand drought and unsuitable for dry farming.
·
Sandy soils are poor store house of plant nutrients
·
Contain low organic matter
·
Leaching of applied nutrients is very high.
·
In sandy soil, few crops can be grown such as potato, groundnut
and cucumbers.
·
Clay particles play a very important role in soil fertility.
·
Clayey soils are difficult to till and require much skill in handling.
When moist clayey soils are exceedingly sticky and when dry, become very hard
and difficult to break.
·
They have fine pores, and are poor in drainage and aeration.
·
They have a high water holding capacity and poor percolation,
which usually results in water logging.
·
They are generally very fertile soils, in respect of plant
nutrient content. Rice, jute, sugarcane can be grown very successfully in these
soils.
·
Loam and Silt loam soils are highly desirable for cultivation
·
Generally, the best agriculture soils are those contain 10 – 20
per cent clay, 5 – 10 per cent organic matter and the rest equally shared by
silt and sand
Soil
Structure
Soil conditions and characteristics such as water
movement, heat transfer, aeration, and porosity are much influenced by
structure. In fact, the important physical changes imposed by the farmer in
ploughing, cultivating, draining, liming, and manuring his land are structural
rather than textural.
Definition of Soil Structure: The arrangement and organization of primary and
secondary particles in a soil mass is known as soil structure.
Soil structure controls the amount of water and air present in soil. Plant
roots and germinating seeds require sufficient air and oxygen for respiration.
Bacterial activities also depend upon the supply of water and air in the soil.
Formation of soil structure: Soil particles may be present either as
single individual grains or as aggregate i.e. group of particles bound together
into granules or compound particles. These granules or compound particles are
known as secondary particles. A majority of particles in a sandy or silty soil
are present as single individual grains while in clayey soil they are present
in granulated condition. The individual particles are usually solid, while the
aggregates are not solid but they possess a porous or spongy character. Most
soils are mixture of single grain and compound particle. Soils, which
predominate with single grains are said to be structure less, while those
possess majority of secondary particles are said to be aggregate, granulated or
crumb structure.
Mechanism of Aggregate Formation: The bonding of the soil particles into
structural unit is the genesis of soil structure. The bonding between
individual particles in the structural units is generally considered to be
stronger than the structural units themselves.
In aggregate formation, a number of primary particles such as sand, silt and
clay are brought together by the cementing or binding effect of soil colloids.
The cementing materials taking part in aggregate formation are colloidal clay,
iron and aluminium hydroxides and decomposing organic matter. Whatever may be
the cementing material, it is ultimately the dehydration of colloidal matter
accompanied with pressure that completes the process of aggregation.
Colloidal clay: By virtue of high surface area and surface charge, clay
particles play a key role in the formation of soil aggregates. Sand and silt
particles can not form aggregates as they do not possess the power of adhesion
and cohesion. These particles usually carry a coating of clay particles; they
are enmeshed in the aggregates formed by the adhering clay particles. Colloidal
particles form aggregates only when they are flocculated. There is vast
difference between flocculation and aggregation.
Flocculation is brought about by coalescence of colloidal particles and is the
first step in aggregation.
Aggregation is some thing more than flocculation involving a combination of
different factors such as hydration, pressure, dehydration etc. and required
cementation of flocculated particles. The cementation may be caused by cations,
oxides of Fe and Al, humus substances and products of microbial excretion and
synthesis. Clay particles form aggregates only if they are wetted by a liquid
like water whose molecules possess an appreciable dipole moment.
Clay - - +Water - - +Cation+ - -Clay - - +Water - -
+Cation+ - -Clay -
The aggregation also depends upon the nature of
clay particles, size and amount of clay particles, dehydration of clay
particles, cations like calcium and anions like phosphate.
Fe and Al oxides: The colloidal Fe oxides act as cementing agent in
aggregation. Al oxides bind the sand and silt particles. These act in two ways.
A part of the hydroxides acts as a flocculating agent and the rest as a
cementing agent.
Organic matter: It also plays an important role in forming soil
aggregates.
- During decomposition, cellulose substances
produce a sticky material very much resembling mucus or mucilage. The
sticky properly may be due to the presence of humic or humic acid or
related compounds produced.
- Certain polysaccharides formed during
decomposition.
- Some fungi and bacteria have cementing effect
probably due to the presence of slimes and gums on the surface of the
living organisms produced as a result of the microbial activity
Classification of
Soil Structure: The primary
particles sand, silt and clay usually occur grouped together in the form of
aggregates.
Natural aggregates are called peds where as clod is
an artificially formed soil mass. Structure is studied in the field under
natural conditions and it is described under three categories
1. Type - Shape or form and arrangement pattern of peds
2. Class - Size of Peds
3. Grade - Degree of distinctness of peds
Types of Soil
Structure: There are four
principal forms of soil structure
Plate-like (Platy): In this type, the aggregates are arranged in
relatively thin horizontal plates or leaflets. The horizontal axis or
dimensions are larger than the vertical axis. When the units/ layers are thick
they are called platy. When they are thin then it is laminar.
Platy structure is most noticeable in the surface layers of virgin soils but
may be present in the subsoil.
This type is inherited from the parent material,
especially by the action of water or ice.
Prism-like: The vertical axis is more developed than
horizontal, giving a pillar like shape. Vary in length from 1- 10 cm. commonly
occur in sub soil horizons of Arid and Semi arid regions. When the tops are
rounded, the structure is termed as columnar when the tops are flat / plane,
level and clear cut prismatic.
Block like: All three dimensions are about the same size. The
aggregates have been reduced to blocks. Irregularly six faced with their three
dimensions more or less equal.
When the faces are flat and distinct and the edges are sharp angular, the
structure is named as angular blocky. When the faces and edges are mainly
rounded it is called sub angular blocky. These types usually are confined to
the sub soil and characteristics have much to do with soil drainage, aeration
and root penetration.
Spheroidal (Sphere
like): All rounded
aggregates (peds) may be placed in this category. Not exceeding an inch in
diameter. These rounded complexes usually loosely arranged and readily
separated. When wetted, the intervening spaces generally are not closed so
readily by swelling as may be the case with a blocky structural condition.
Therefore in sphere like structure, infiltration,
percolation and aeration are not affected by wetting of soil.
The aggregates of this group are usually termed as granular which are
relatively less porous. When the granules are very porous, it is termed as
crumb. This is specific to surface soil particularly high in organic matter/
grass land soils.
Classes of Soil
Structure: Each primary
structural type of soil is differentiated into 5 size classes depending upon
the size of the individual peds.
The terms commonly used for the size classes are:
- 1. Very fine or very thin
- 2. Fine or thin
- 3. Medium
- 4. Coarse or thick
- 5. Very Coarse or very thick
The terms thin and thick are used for platy types,
while the terms fine and coarse are used for other structural types.
Grades of Soil
Structure: Grades indicate
the degree of distinctness of the individual peds. It is determined by the stability
of the aggregates. Grade of structure is influenced by the moisture content of
the soil. Grade also depends on organic matter, texture etc. Four terms
commonly used to describe the grade of soil structure are:
- Structure less: There is no noticeable
aggregation, such as conditions exhibited by loose sand.
- Weak Structure: Poorly formed, indistinct
formation of peds, which are not durable and much unaggregated material.
- Moderate structure: Moderately well developed
peds, which are fairly durable and distinct.
- Strong structure: Very well formed peds, which
are quite durable and distinct.
Soil Structure
Naming: For naming a soil
structure the sequence followed is grade, class and type; for example strong
coarse angular blocky, moderate thin platy, weak fine prismatic.
Factors
Affecting Soil Structure
The development of structure in arable soil depends
on the following factors:
1. Climate: Climate has considerable influence on the degree of aggregation as
well as on the type of structure. In arid regions there is very little
aggregation of primary particles. In semi arid regions, the degree of
aggregation is greater.
2. Organic matter: Organic matter improves the structure of a sandy soil as
well as of a clay soil. In case of a sandy soil, the sticky and slimy material
produced by the decomposing organic matter and the associated microorganism
cement the sand particles together to form aggregates. In case of clayey soil,
it modifies the properties of clay by reducing its cohesiveness. This helps
making clay more crumby.
3. Tillage: Cultivation implements break down the
large clods into smaller fragments and aggregates. For obtaining good granular
and crumby structure, optimum moisture content in the soil is necessary. If the
moisture content is too high it will form large clods on drying. If it is too
low some of the existing aggregates will be broken down.
4. Plants, Roots and Residues: Excretion of
gelatinous organic compounds and exudates from roots serve as a link. Root
hairs make soil particles to cling together. Grass and cereal roots vs other
roots. Pressure exerted by the roots also held the particles together.
Dehydration of soil strains the soil due to shrinkage result in cracks lead to
aggregation
Plant tops and residues shade the soil prevent it from extreme and sudden
temperature and moisture changes and also from rain drop impedance.
Plant residues serve as a food to microbes which are the prime aggregate
builders.
5. Animals: Among the soil fauna small animals like earthworms, moles and
insects etc., that burrow in the soil are the chief agents that take part in
the aggregation of finer particles.
6. Microbes: Algae, fungi, actinomycetes and fungi
keep the soil particles together. Fungi and actinomycetes exert mechanical
binding by mycelia, Cementation by the products of decomposition and materials
synthesized by bacteria.
7. Fertilizers: Fertilizer like Sodium Nitrate destroys granulation by reducing
the stability of aggregates. Few fertilizers for example, CAN help in
development of good structures.
8. Wetting and drying: When a dry soil is wetted,
the soil colloids swell on absorbing water. On drying, shrinkage produces
strains in the soil mass gives rise to cracks, which break it up into clods and
granules of various sizes.
9. Exchangeable cations: Ca, Mg
----------à
H, Na
Flocculating
Deflocculating
Good structure Poor
structure
10. Inorganic
cements: CaCO3 and
Sesquioxides
11. Clay
12. Water
Effect of
Soil Structure on other Physical Properties
Porosity: Porosity of a soil is easily changed. In
plate like structure, pore spaces are less where as in crumby structure pore
spaces are more.
Temperature: Crumby structure provides good
aeration and percolation of water in the soil. Thus these characteristics help
in keeping optimum temperature in comparison to plate like structure.
Density: Bulk density varies with the total pore
space present in the soil. Structure chiefly influences pore spaces Platy
structure with less total pore spaces has high bulk density where as crumby
structure with more total pore spaces has low bulk density.
Consistence: Consistence of soil also depends on
structure. Plate-like structure exhibits strong plasticity.
Colour: Bluish and greenish colors of soil are
generally due to poor drainage of soil. Platy structure normally hinders free
drainage.
Importance of Structure: Soil structure influences
rather indirectly by the formation of an array of pores of various shapes and
sizes. These pores are controlling factors governing water, air and temperature
in soil.
Role of
soil structure in relation to plant growth
- Soil structure influences the amount and
nature of porosity.
- Structure controls the amount of water and air
present in the soil. Not only the amount of water and air dependent on
soil structure, but their movement and circulation are also controlled by
soil structure.
- It affects tillage practices.
- Structure controls runoff and erosion.
- Platy structure normally hinders free drainage
whereas sphere like structure (granular and crumby) helps in drainage.
- Crumby and granular structure provides optimum
infiltration, water holding capacity, aeration and drainage. It also
provides good habitat for microorganisms and supply of nutrients.
Class of Soil Structure as differentiated by size
of soil peds
|
Class
|
Platy
|
Prism-atic
|
Colum-nar
|
Blocky
|
S.A. Blocky
|
Gran-ular
|
Crumb
|
|
V.Fine or V.Thin
|
<1
|
<10
|
<10
|
<5
|
<5
|
<1
|
<1
|
|
Fine or Thin
|
1-2
|
10-20
|
10-20
|
5-10
|
5-10
|
1-2
|
1-2
|
|
Medi
|
2-5
|
20-50
|
20-50
|
10-20
|
10-20
|
2-5
|
2-5
|
|
um
|
5-10
|
50-100
|
50-100
|
20-50
|
20-50
|
5-10
|
-
|
|
Coarseor Thick
|
>10
|
>100
|
>100
|
>50
|
>50
|
>10
|
-
|
|
V.C
|
>10
|
>100
|
>100
|
>50
|
>50
|
>10
|
-
|
|
or V.Thic
|
Role of
soil structure in relation to plant growth
- Soil structure influences the amount and
nature of porosity.
- Structure controls the amount of water and air
present in the soil. Not only the amount of water and air dependent on
soil structure, but their movement and circulation are also controlled by
soil structure.
- It affects tillage practices.
- Structure controls runoff and erosion.
- Platy structure normally hinders free drainage
whereas sphere like structure (granular and crumby) helps in drainage.
- Crumby and granular structure provides optimum
infiltration, water holding capacity, aeration and drainage. It also
provides good habitat for microorganisms and supply of nutrients.
Class of Soil Structure as differentiated by size
of soil peds
|
Class
|
Platy
|
Prism-atic
|
Colum-nar
|
Blocky
|
S.A. Blocky
|
Gran-ular
|
Crumb
|
|
V.Fine or V.Thin
|
<1
|
<10
|
<10
|
<5
|
<5
|
<1
|
<1
|
|
Fine or Thin
|
1-2
|
10-20
|
10-20
|
5-10
|
5-10
|
1-2
|
1-2
|
|
Medi
|
2-5
|
20-50
|
20-50
|
10-20
|
10-20
|
2-5
|
2-5
|
|
um
|
5-10
|
50-100
|
50-100
|
20-50
|
20-50
|
5-10
|
-
|
|
Coarseor Thick
|
>10
|
>100
|
>100
|
>50
|
>50
|
>10
|
-
|
|
V.C
|
>10
|
>100
|
>100
|
>50
|
>50
|
>10
|
-
|
|
or V.Thic
|
Density of
Soil: Bulk Density and Particle Density
Density represents weight (mass) per unit volume of
a substance.
Density = Mass / Volume
Soil density is expressed in two well accepted
concepts as particle density and bulk density. In the metric system, particle
density can be expressed in terms of mega grams per cubic meter (Mg/m3). Thus
if 1 m3 of soil solids weighs 2.6 Mg, the particle density is 2.6 Mg / m3
(since 1 Mg =1 million grams and 1 m3 =1 million cubic centimeters) thus
particle density can also be expressed as 2.6 g / cm3.
Particle Density: The weight per unit volume of the
solid portion of soil is called particle density. Generally particle density of
normal soils is 2.65 grams per cubic centimeter. The particle density is higher
if large amount of heavy minerals such as magnetite; limonite and hematite are
present in the soil. With increase in organic matter of the soil the particle
density decreases. Particle density is also termed as true density.
Table Particle density of different soil textural
classes
|
Textural classes
|
Particle
density ( g/ cm3)
|
|
Coarse sand
|
2.655
|
|
Fine sand
|
2.659
|
|
Silt
|
2.798
|
|
Clay
|
2.837
|
Bulk Density: The oven dry weight of a unit volume
of soil inclusive of pore spaces is called bulk density. The bulk density of a
soil is always smaller than its particle density. The bulk density of sandy
soil is about 1.6 g / cm3, whereas that of organic matter is about 0.5. Bulk
density normally decreases, as mineral soils become finer in texture. The bulk
density varies indirectly with the total pore space present in the soil and
gives a good estimate of the porosity of the soil. Bulk density is of greater
importance than particle density in understanding the physical behavior of the
soil. Generally soils with low bulk densities have favorable physical
conditions.
Bulk density of different textural classes
|
Textural class
|
Bulk
density (g/cc)
|
Pore space
(%)
|
|
Sandy soil
|
1.6
|
40
|
|
Loam
|
1.4
|
47
|
|
Silt loam
|
1.3
|
50
|
|
Clay
|
1.1
|
58
|
Factors affecting bulk density
1. Pore space: Since bulk density relates to the combined volume of the solids
and pore spaces, soils with high proportion of pore space to solids have lower
bulk densities than those that are more compact and have less pore space.
Consequently, any factor that influences soil pore space will affect bulk
density.
2. Texture: Fine textured surface soils such as
silt loams, clays and clay loams generally have lower bulk densities than sandy
soils. This is because the fine textured soils tend to organize in porous
grains especially because of adequate organic matter content. This results in
high pore space and low bulk density. However, in sandy soils, organic matter
content is generally low, the solid particles lie close together and the bulk
density is commonly higher than in fine textured soils.
3. Organic matter content: More the organic matter content in soil results in
high pore space there by shows lower bulk density of soil and vice-versa.
Soil
Colloids
The colloidal state refers to a two-phase system in
which one material in a very finely divided state is dispersed through second
phase.
The examples are:
Solid in liquid - Clay in water (dispersion of clay
in water)
Liquid in gas -Fog or clouds in atmosphere
The clay fraction of the soil contains particles
less than 0.002 mm in size. Particles less than 0.001 mm size possess colloidal
properties and are known as soil colloids.
General Properties of Soil Colloids
1. Size: The most important common property of
inorganic and organic colloids is their extremely small size. They are too
small to be seen with an ordinary light microscope. Only with an electron
microscope they can be seen. Most are smaller than 2 micrometers in diameter.
2. Surface area: Because of their small size, all
soil colloids expose a large external surface per unit mass. The external
surface area of 1 g of colloidal clay is at least 1000 times that of 1 g of
coarse sand. Some colloids, especially certain silicate clays have extensive
internal surfaces as well. These internal surfaces occur between plate like
crystal units that make up each particle and often greatly exceed the external
surface area. The total surface area of soil colloids ranges from 10 m2/g for
clays with only external surfaces to more than 800 m2/g for clays with
extensive internal surfaces. The colloid surface area in the upper 15 cm of a
hectare of a clay soil could be as high 700,000 km2/g
3. Surface charges: Soil colloidal surfaces, both
external and internal characteristically carry negative and/or positive
charges. For most soil colloids, electro negative charges predominate. Soil
colloids both organic and inorganic when suspended in water, carry a negative
electric charge. When an electric current is passed through a suspension of
soil colloidal particles they migrate to anode, the positive electrode
indicating that they carry a negative charge. The magnitude of the charge is
known as zeta potential. The presence and intensity of the particle charge
influence the attraction and repulsion of the particles towards each other,
there by influencing both physical and chemical properties.
The negative electrical charge on clays comes from
i) Ionizable hydrogen ions and
ii) Isomorphism substitution.
i) Ionizable hydrogen ions: Ionizable hydrogen ions are hydrogen from hydroxyl
ions on clay surfaces. The -Al-OH or -Si-OH portion of the clay ionizes the H
and leaves an unneutralized negative charge on the oxygen (-Al-O- or - Si-O).
The extent of ionized hydrogen depends on solution pH; more ionization occurs in
more alkaline (basic) solutions.
ii) Isomorphous substitution: The second source of charge on clay particles is
due to the substitution of one ion for another of similar size and often with
lower positive valence. In clay structures, certain ions fit into certain
mineral lattice sites because of their convenient size and charge. Dominantly,
clays have Si4+ in tetrahedral sites and A13+ in octahedral sites. Other ions
present in large amounts during clay crystallization can replace some of the
A13+ and Si4+ cations. Substitutions that are common are the Si4+ replaced by
A13+, and even more extensive replacement of A13+ by one or more of these:
Fe3+, Fe2+, Mg2+ or Zn2+ Since the total negative charge from the anions (the
oxygen) remains unchanged, the lower positive charge because of substitution
results in an excess negative charge at that location in the structure.
4. Adsorption of cations: As soil colloids possess
negative charge they attract the ions of an opposite charge to the colloidal
surfaces. They attract hundreds of positively charged ions or cation such as
H+, A13+ Ca2+ , and Mg2+. This gives rise to an ionic double layer.
The process, called Isomorphous substitution and
the colloidal particle constitutes the inner ionic layer, being essentially huge
anions; with both, external and internal layers that are negative in charge.
The outer layer is made up of a swarm of rather loosely held (adsorbed) cations
attracted to the negatively charged surfaces. Thus a colloidal particle is
accompanied by a swarm of cations that are adsorbed or held on the particle
surfaces.
5. Adsorption of water: In addition to the adsorbed
cations, a large number of water molecules are associated with soil colloidal
particles. Some are attracted to the adsorbed cations, each of which is
hydrated; others are held in the internal surfaces of the colloidal particles.
These water molecules play a critical role in determining both the physical and
chemical properties of soil.
5. Cohesion: Cohesion is the phenomenon of sticking
together of colloidal particles that are of similar nature. Cohesion indicates
the tendency of clay particles to stick together. This tendency is primarily
due to the attraction of the clay particles for the water molecules held
between them. When colloidal substances are wetted, water first adheres to the
particles and then brings about cohesion between two or more adjacent colloidal
particles.
6. Adhesion: Adhesion refers to the phenomenon of
colloidal particles sticking to other substances. It is the sticking of
colloida1 materials to the surface of any other body or substance with which it
comes in contact.
7. Swelling and shrinkage: Some clay (soil
colloids) such as smectites swell when wet and shrink when dry. After a
prolonged dry spell, soils high in smectites (e.g. Vertisols) often are
crises-crossed by wide, deep cracks, which at first allow rain to penetrate
rapidly. Later, because of swelling, such soil is likely to close up and become
much more impervious than one dominated by kaolinite, chlorite, or fine grained
micas. Vermiculite is intermediate in its swelling and shrinking
characteristics.
8. Dispersion and flocculation: As long as the
colloidal particles remain charged, they repel each other and the suspension
remains stable. If on any account they loose their charge, or if the magnitude
of the charge is reduced, the particles coalesce, form flocs or loose
aggregates, and settle out. This phenomenon of coalescence and formation of
flocs is known as flocculation. The reverse process of the breaking up of flocs
into individual particles is known as deflocculation or dispersion.
9. Brownian movement: When a suspension of
colloidal particles is examined under a microscope the particles seem to
oscillate. The oscillation is due to the collision of colloidal particles or
molecules with those of the liquid in which they are suspended. Soil colloidal
particles with those of water in which they are suspended are always in a
constant state of motion. The smaller the particle, the more rapid is its movement.
10. Non permeability: Colloids, as opposed to
crystalloid, are unable to pass through a semi-permeable membrane. Even though
the colloidal particles are extremely small, they are bigger than molecules of
crystalloid dissolved in water. The membrane allows the passage of water and of
the dissolved substance through its pores, but retains the colloidal particles.
Types of Soil
Colloids
There
are four major types of colloids present in soil
1. Layer silicate clays
2. Iron and aluminum oxide clays (sesquioxide clays)
3. Allophane and associated amorphous clays
4. Humus.
Layer
silicate clays, iron and aluminum oxide clays, allophane and associated
amorphous clays are inorganic colloids while humus is an organic colloid.
1. Layer silicate clays: These are
most important silicate clays and are known as phyllosilicates (Phyllon - leaf)
because of their leaf-like or plate like structure. They are comprised of two
kinds of horizontal sheets. One dominated by silicon and other by aluminum
and/or magnesium.
Silica Tetrahedron: The basic
building block for the silica-dominated sheet is a unit composed of one silicon
atom surrounded by four oxygen atoms. It is called the silica tetrahedron
because of its four-sided configuration. An interlocking array or a series of
these silica tetrahedral tied together horizontally by shared oxygen anions
gives a tetrahedral sheet.
Alumina
Octahedron: Aluminium and/or magnesium ions are the key cations
in the second type of sheet. An aluminium (or magnesium) ion surrounded by six
oxygen atoms or hydroxyl group gives an eight sided building block termed
octahedron. Numerous octahedra linked together horizontally comprise the
octahedral sheet. An aluminum dominated sheet is known as a dioctahedral sheet,
whereas one dominated by magnesium is called a trioctahedral sheet. The
distinction is due to the fact that two aluminum ions in a dioctahedral sheet
satisfy the same negative charge from surrounding oxygen and hydroxyls as three
magnesium ions in a trioctahedral sheet.
The
tetrahedral and octahedral sheets are the fundamental structural units of
silicate clays. They, in turn, are bound together within the crystals by shared
oxygen atoms into different layers. The specific nature and combination of
sheets in these layers vary from one type of clay to another and largely
control the physical and chemical properties of each clay.
Types
of Silicate Clay Minerals: On the basis of the number and
arrangement of tetrahedral (silica) and octahedral (alumina-magnesia) sheets
contained in the crystal units or layers, silicate clays are classified into
three different groups
A.
1 :1 Type clay minerals
B.
2:1 Type clay minerals
C.
2: 1: 1 Type clay minerals
A. 1:1 Type Minerals: The layers of the 1:1-type minerals are made up of one
tetrahedral (silica) sheet combined with one octahedral (alumina) sheet-hence
the terminology. In soils, kaolinite is the most prominent member of this
group, which includes hallosite, nacrite, and dickite.
The tetrahedral and octahedral sheets in a layer of a kaolinite
crystal are held together tightly by oxygen atoms, which are mutually shared by
the silicon and aluminum cations in their respective sheets. These layers, in
turn, are held together by hydrogen bonding. Consequently, the structure is
fixed and no expansion ordinarily occurs between layers when the clay is
wetted.
Cations and water do not enter between the structural layers of a
1:1 type mineral particle. The effective surface of kaolinite is thus
restricted to its outer faces or to its external surface area. Also, there is
little isomorphous substitution in this 1:1 type mineral. Along with the
relatively low surface area of kaolinite, this accounts for its low capacity to
adsorb cations.
Kaolinite crystals usually are hexagonal in shape. In comparison
with other clay particles, they are large in size, ranging from 0.10 to 5 um
across with the majority falling within the 0.2 to 2 um range. Because of the
strong binding forces between their structural layers, kaolinite particles are
not readily broken down into extremely thin plates.
Kaolinite exhibits very little plasticity (capability of being molded),
cohesion, shrinkage, and swelling.
B. 2:1-Type Minerals: The crystal units (layers) of these minerals are characterized by
an octahedral sheet sandwiched between two tetrahedral sheets. Three general
groups have this basic crystal structure.
i) Expanding type: smectites and vermiculite
ii) Non-expanding type: mica (illite)
Expanding Minerals: The smectite group is noted for interlayer expansion, which
occurs by swelling when the minerals are wetted, the water entering the
interlayer space and forcing the layers apart. Montmorillonite is the most
prominent member of this group in soils, although beidellite, nontronite, and
saponite are also found.
The flake-like crystals of smectites (e.g., Montmorillonite) are
composed of an expanding lattice 2:1-type clay mineral. Each layer is made up
of an octahedral sheet sandwiched between two tetrahedral (silica) sheets.
There is little attraction between oxygen atoms in the bottom tetrahedral sheet
of one unit and those in the top tetrahedral sheet of another. This permits a
ready and variable space between layers, which is occupied by water and exchangeable
cations. This internal surface far exceeds the surface around the outside of
the crystal. In montmorillonite magnesium has replaced aluminum in some sites
of the octahedral sheet. Likewise, some silicon atoms in the tetrahedral sheet
may be replaced by aluminum. These substitutions give rise to a negative
charge.
These minerals show high cation exchange capacity, marked swelling
and shrinkage properties. Wide cracks commonly form as smectite dominated soils
(e.g., Vertisols) are dried. The dry aggregates or clods are very hard, making
such soils difficult to till.
Vermiculites are also 2: 1 type minerals in that an octahedral
sheet occurs between two tetrahedral sheets. In most soils vermiculites, the
octahedral sheet is aluminum dominated (dioctahedral), although magnesium
dominated (trioctahedral) vermiculites are also common. In the tetrahedral
sheet of most vermiculite, considerable substitution of aluminum for silicon
has taken place. This accounts for most of the very high net negative charge associated
with these minerals.
Water molecules, along with magnesium and other ions, are strongly adsorbed in
the interlayer space of vermiculites. They act primarily as bridges holding the
units together rather than as wedges driving them apart. The degree of swelling
is, therefore considerable less for vermiculites than for smectites. For this
reason, vermiculites are considered limited-expansion clay minerals, expanding
more than kaolinite but much less than the smectites.
The cation exchange capacity of vermiculites usually exceeds that
of all other silicate clays, including montmorillonite and other smectites,
because of very high negative charge in the tetrahedral sheet. Vermiculite
crystals are larger than those of the smectites but much smaller than those of
kaolinite.
Non-expanding minerals: Micas are the type minerals in this group.
(e.g.) Muscovite and biotite, weathered minerals similar in structure to these
micas are found in the clay fraction of soils. They are called fine-grained
micas. Like sanctities, fine-grained micas have a 2:1-type crystal. However,
the particles are much larger than those of the smectites. Also, the major
source of charge is in-the tetrahedral sheet where aluminum atoms occupy about
20% of the silicon sites. These results in a high net negative charge in the
tetrahedral sheet, even higher than that found in vermiculites, to satisfy this
charge, potassium ions are strongly attracted in the interlayer space and are
just the right size to fit into certain spaces in the adjoining tetrahedral
sheets. The potassium thereby acts as a binding agent, preventing expansion of
the crystal. Hence, fine-grained micas are quite non expansive.
The properties such as hydration, cation adsorption, swelling,
shrinkage, and plasticity are much less intense in fine-grained micas than in
smectites. The fine grained micas exceed kaolinite with respect to these
characteristics, but this may be due in part to the presence of interstratified
layers of smectite or vermiculite. In size, too, fine-grained mica crystals are
intermediate between the smectities and kaolihites.
Their specific surface area varies from 70 to 100 m2/g, about one eighth that
for the smectites.
C. 2:1:1 Type Minerals: This silicate group is represented by chlorites, which are common
in a variety of soils. Chlorites are basically iron magnesium silicates with
some aluminum present. In a typical chlorite clay crystal, 2:1 layers, such as
in vermiculites, alternate with a magnesium-dominated trioctahedral sheet,
giving a 2:1:1 ratio. Magnesium also dominates the trioctahedral sheet in the
2:1 layer of chlorites. Thus, the crystal unit contains two silica tetrahedral
sheets and two magnesium-dominated trioctahedral sheets giving rise to the term
2:1:1 or 2:2-type structure.
The negative charge of chlorites is about the same as that of fine-grained
micas considerably less than that of the smectites or vermiculites. Like fine
micas, chlorites may be interstratified with vermiculites or smectites in a
single crystal. Particle size and surface area for chlorites are also about the
same as for fine grained micas. There is no water adsorption between the
chlorite crystal units, which accounts for the non expanding nature of this
mineral.
Mixed and interstratified layers: Specific
groups of clay minerals do not occur independently of one another. In a given
soil, it is common to find several clay minerals in an intimate mixture.
Furthermore, some mineral colloids have properties and composition intermediate
between those of any two of the well defined minerals described. Such minerals
are termed mixed layer or interstratified because the individual layers within
a given crystal may be of more than one type. Terms such as
"chlorite-vermiculite" and "fine-grained mica- smectite"
are used to describe mixed-layer minerals. In some soils, they are more common
than single-structured minerals such as montmorillonite.
2. Iron and aluminum oxide clays (sesquioxide clays):
Under conditions of extensive leaching by rainfall and long time intensive weathering
of minerals in humid warm climates, most of the silica and much of the alumina
in primary minerals are dissolved and slowly leached away. The remnant
materials, which have lower solubility, are sesquioxides. Sesquioxides (metal
oxides) are mixtures of aluminum hydroxide, Al (OH)3, and iron oxide, Fe2O3, or
iron hydroxide, Fe (OH)3.
The Latin word sesqui means one and one- half times, meaning one and one-half
times more oxygen than Al and Fe. These clays can grade from amorphous to
crystalline.
Examples
of iron and aluminum oxides common in soils are gibbsite (Al2O3.3H2O) and
geothite (Fe2O3.H2O).
Less
is known about these clays than about the layer silicates. These clays do not
swell, not sticky and have high phosphorus adsorption capacity
3. Allophane and other Amorphous Minerals: These
silicate clays are mixtures of silica and alumina. They are amorphous in
nature. Even mixture of other weathered oxides (iron oxide) may be a part of
the mixture. Typically, these clays occur where large amount of weathered
products existed. These clays are common in soils forming from volcanic ash
(e.g., Allophane). These clays have high anion exchange capacity or even high
cation exchange capacity. Almost all of their charge is from accessible
hydroxyl ions (OH-), which can attract a positive ion or lose the H+ attached.
These clays have a variable charge that depends on H+ in solution (the soil
acidity).
4. Humus (Organic Colloid): Humus is
amorphous, dark brown to black, nearly insoluble in water, but mostly soluble
in dilute alkali (NaOH or KOH) solutions. It is a temporary intermediate
product left after considerable decomposition of plant and animal remains. They
are temporary intermediate because the organic substances remain continue to
decompose slowly.
The
humus is often referred to as an organic colloid and consists of various chains
and loops of linked carbon atoms. The humus colloids are not crystalline. They
are composed basically of carbon, hydrogen, and oxygen rather than of silicon,
aluminum, iron, oxygen, and hydroxyl groups.
The
organic colloidal particles vary in size, but they may be at least as small as
the silicate clay particles. The negative charges of humus are associated with
partially dissociated enolic (-OH), carboxyl (-COOH), and phenolic groups;
these groups in turn are associated with central units of varying size and
complexity.
Soil
Moisture
Water contained in soil is called soil moisture.
The water is held within the soil pores. Soil water is the major component of
the soil in relation to plant growth. If the moisture content of a soil is
optimum for plant growth, plants can readily absorb soil water. Not all the
water, held in soil, is available to plants. Much of water remains in the soil
as a thin film. Soil water dissolves salts and makes up the soil solution,
which is important as medium for supply of nutrients to growing plants.
Importance of Soil Water
- Soil water serves as a solvent and carrier of
food nutrients for plant growth
- Yield of crop is more often determined by the
amount of water available rather than the deficiency of other food
nutrients
- Soil water acts as a nutrient itself
- Soil water regulates soil temperature
- Soil forming processes and weathering depend
on water
- Microorganisms require water for their
metabolic activities
- Soil water helps in chemical and biological
activities of soil
- It is a principal constituent of the growing
plant
- Water is essential for photosynthesis
Retention of Water by Soil: The soils hold water
(moisture) due to their colloidal properties and aggregation qualities. The
water is held on the surface of the colloids and other particles and in the
pores. The forces responsible for retention of water in the soil after the
drainage has stopped are due to surface tension and surface attraction and are
called surface moisture tension. This refers to the energy concept in moisture
retention relationships. The force with which water is held is also termed as
suction.
Ways to
Retain Water in Soil
1. Cohesion and adhesion forces: These two basic
forces are responsible for water retention in the soil. One is the attraction
of molecules for each other i.e., cohesion. The other is the attraction of
water molecules for the solid surface of soil i.e. adhesion. By adhesion,
solids (soil) hold water molecules rigidly at their soil - water interfaces.
These water molecules in turn hold by cohesion. Together, these forces make it
possible for the soil solids to retain water.
2. Surface tension: This phenomenon is commonly
evidenced at water- air interfaces. Water behaves as if its surface is covered
with a stretched elastic membrane. At the surface, the attraction of the air
for the water molecules is much less than that of water molecules for each
other. Consequently, there is a net downward force on the surface molecules,
resulting in sort of a compressed film (membrane) at the surface. This
phenomenon is called surface tension.
3. Polarity or dipole character: The retention of
water molecules on the surface of clay micelle is based on the dipole character
of the molecule of water. The water molecules are held by electrostatic force
that exists on the surface of colloidal particles. By virtue of their dipole
character and under the influence of electrostatic forces, the molecules of
water get oriented (arranged) on the surface of the clay particles in a
particular manner.
Each water molecule carries both negative and positive
charges. The clay particle is negatively charged. The positive end of water
molecule gets attached to the negatively charged surface of clay and leaving
its negative end outward. The water molecules attached to the clay surface in
this way present a layer of negative charges to which another layer of oriented
water molecules is attached. The number of successive molecular layers goes on
increasing as long as the water molecules oriented. As the molecular layer gets
thicker, orientation becomes weaker, and at a certain distance from the
particle surface the water molecules cease to orientate and capillary water
(liquid water) begins to appear. Due to the forces of adsorption (attraction)
exerted by the surface of soil particles, water gets attached on the soil
surface. The force of gravity also acts simultaneously, which tries to pull it
downwards. The surface force is far greater than the force of gravity so water
may remain attached to the soil particle. The water remains attached to the
soil particle or move downward into the lower layers, depending on the
magnitude of the resultant force.
Factors Affecting Soil Water
1. Texture: Finer the texture, more is the pore space and also
surface area, greater is the retention of water.
2. Structure: Well-aggregated porous structure favors better
porosity, which in turn enhance water retention.
3. Organic matter: Higher the organic matter more is the water
retention in the soil.
4. Density of soil: Higher the density of soil, lower is the moisture
content.
5. Temperature: Cooler the temperature, higher is the moisture
retention.
6. Salt content: More the salt content in the soil less is the
water available to the plant.
7. Depth of soil: More the depth of soil more is the water available
to the plant.
8. Type of clay: The 2:1 type of day increases the water retention
in the soil.
Soil Water
Potential
The retention and movement of water in soils, its
uptake and translocation in plants and its loss to the atmosphere are all
energy related phenomenon. The more strongly water is held in the soil the
greater is the heat (energy) required. In other words, if water is to be
removed from a moist soil, work has to be done against adsorptive forces.
Conversely, when water is adsorbed by the soil, a negative amount of work is
done. The movement is from a zone where the free energy of water is high (standing
water table} to one where the free energy is low (a dry soil). This is called
soil water energy concept.
Free energy of soil solids for water is affected by
i) Matric (solid) force i.e., the attraction of the
soil solids for water (adsorption} which markedly reduces the free energy
(movement} of the adsorbed water molecules.
ii) Osmotic force i.e., the attraction of ions and
other solutes for water to reduce the free energy of soil solution.
Matric and Osmotic potentials are negative and
reduce the free energy level of the soil water. These negative potentials are
referred as suction or tension.
iii) Force of gravity: This acts on soil water, the
attraction is towards the earth's center, which tends to pull the water down
ward. This force is always positive. The difference between the energy states
of soil water and pure free water is known as soil water potential. Total water
potential (Pt} is the sum of the contributions of gravitational potential (Pg),
matric potential (Pm) and the Osmotic potential or solute potential (Po).
Pt = Pg + Pm + Po
Potential represents the difference in free energy
levels of pure water and of soil water. The soil water is affected by the force
of gravity, presence of soil solid (matric) and of solutes.
Methods of expressing suctions: There are two units
to express differences in energy levels of soil water.
i) PF Scale: The free energy is measured in terms
of the height of a column of water required to produce necessary suction or
pressure difference at a particular soil moisture level. The pF, therefore,
represents the logarithm of the height of water column (cm) to give the
necessary suction.
ii) Atmospheres or Bars: It is another common mean
of expressing suction. Atmosphere is the average air pressure at sea level. If
the suction is very low as occurs in the case of a wet soil containing the
maximum amount of water that it can hold, the pressure difference is of the
order of about 0.01 atmospheres or 1 PF equivalent to a column of water 10 cm
in height. Similarly, if the pressure difference is 0.1 atmosphere the PF will
be 20. Soil moisture constants can be expressed in term of PF values. A soil
that is saturated with water has PF 0 while an oven dry soil has a PF 7.
Measuring
Soil Moisture
Two general types of measurements relating to soil
water are ordinarily used
i) By some methods the moisture content is measured directly or indirectly
ii) Techniques are used to determine the soil moisture potential (tension or
suction)
Measuring soil moisture content in laboratory
1. Gravimetric method: This consists of obtaining a
moist sample, drying it in an oven at 105°C until it losses no more weight and
then determining the percentage of moisture. The gravimetric method is time consuming
and involves laborious processes of sampling, weighing and drying in
laboratory.
2. Electrical conductivity method: This method is
based upon the changes in electrical conductivity with changes in soil
moisture. Gypsum blocks inside of with two electrodes at a definite distance
are apart used in this method. These blocks require previous calibration for
uniformity. The blocks are buried in the soil at desired depths and the
conductivity across the electrodes measured with a modified Wheatstone bridge.
These electrical measurements are affected by salt concentration in the soil
solution and are not very helpful in soils with high salt contents.
Measuring soil moisture potential in situ (field)
3. Suction method or equilibrium tension method: Field tensiometers measure the
tension with which water is held in the soils. They are used in determining the
need for irrigation. The tensiometer is a porous cup attached to a glass tube,
which is connected to a mercury monometer. The tube and cup are filled with
water and cup inserted in the soil. The water flows through the porous cup into
the soil until equilibrium is established. These tension readings in monometer,
expressed in terms of cm or atmosphere, measures the tension or suction of the
soil.
If the soil is dry, water moves through the porous cup, setting up a negative
tension (or greater is the suction). The tensiometers are more useful in sandy
soils than in fine textured soils. Once the air gets entrapped in the
tensiometer, the reliability of readings is questionable.
Classification
of soil water
Soil water has been classified from a physical and
biological point of view as Physical classification of soil water, and
biological classification of soil water.
Physical classification of soil water
1. Gravitational water: Gravitational water occupies the larger soil pores
(macro pores) and moves down readily under the force of gravity. Water in
excess of the field capacity is termed gravitational water. Gravitational water
is of no use to plants because it occupies the larger pores. It reduces
aeration in the soil. Thus, its removal from soil is a requisite for optimum
plant growth. Soil moisture tension at gravitational state is zero or less than
1/3 atmosphere.
Factors affecting gravitational water
i. Texture: Plays a great role in controlling the
rate of movement of gravitational water. The flow of water is proportional to
the size of particles. The bigger the particle, the more rapid is the flow or
movement. Because of the larger size of pore, water percolates more easily and
rapidly in sandy soils than in clay soils.
ii. Structure: It also affects gravitational water.
In platy structure movement of gravitational water is slow and water stagnates
in the soil. Granular and crumby structure helps to improve gravitational water
movement. In clay soils having single grain structure, the gravitational water,
percolates more slowly. If clay soils form aggregates (granular structure), the
movement of gravitational water improves.
2. Capillary water: Capillary water is held in the
capillary pores (micro pores). Capillary water is retained on the soil
particles by surface forces. It is held so strongly that gravity cannot remove
it from the soil particles. The molecules of capillary water are free and mobile
and are present in a liquid state. Due to this reason, it evaporates easily at
ordinary temperature though it is held firmly by the soil particle; plant roots
are able to absorb it. Capillary water is, therefore, known as available water.
The capillary water is held between 1/3 and 31 atmosphere pressure.
Factors affecting capillary water: The amount of
capillary water that a soil is able to hold varies considerably. The following
factors are responsible for variation in the amount of capillary water.
i. Surface tension: An increase in surface tension
increases the amount of capillary water.
ii. Soil texture: The finer the texture of a soil,
greater is the amount of capillary water holds. This is mainly due to the
greater surface area and a greater number of micro pores.
iii. Soil structure: Platy structure contains more
water than granular structure.
iv. Organic matter: The presence of organic matter
helps to increase the capillary capacity of a soil. Organic matter itself has a
great capillary capacity. Undecomposed organic matter is generally porous
having a large surface area, which helps to hold more capillary water. The
humus that is formed on decomposition has a great capacity for absorbing and
holding water. Hence the presence of organic matter in soil increases the
amount of capillary water in soil.
3. Hygroscopic water: The water that held tightly
on the surface of soil colloidal particle is known as hygroscopic water. It is
essentially non-liquid and moves primarily in the vapour form.
Hygroscopic water held so tenaciously (31 to 10000
atmospheres) by soil particles that plants can not absorb it. Some
microorganism may utilize hygroscopic water. As hygroscopic water is held
tenaciously by surface forces its removal from the soil requires a certain
amount of energy. Unlike capillary water which evaporates easily at atmospheric
temperature, hygroscopic water cannot be separated from the soil unless it is
heated.
Factors affecting hygroscopic water: Hygroscopic water is held on the surface
of colloidal particles by the dipole orientation of water molecules. The amount
of hygroscopic water varies inversely with the size of soil particles. The
smaller the particle, the greater is the amount of hygroscopic water it
adsorbs. Fine textured soils like clay contain more hygroscopic water than
coarse textured soils.
The amount of clay and also its nature influences
the amount of hygroscopic water. Clay minerals of the montmoril1onite type with
their large surface area adsorb more water than those of the kaolinite type,
while illite minerals are intermediate
Biological
Classification of Soil Water
There is a definite relationship between moisture
retention and its utilization by plants. This classification based on the
availability of water to the plant. Soil moisture can be divided into three
parts.
i. Available water: The water which lies between wilting coefficient and field
capacity. It is obtained by subtracting wilting coefficient from moisture
equivalent.
ii. Unavailable water: This includes the whole of
the hygroscopic water plus a part of the capillary water below the wilting
point.
iii. Super available or superfluous water: The
water beyond the field capacity stage is said to be super available. It
includes gravitational water plus a part of the capillary water removed from
larger interstices. This water is unavailable for the use of plants. The
presence of super-available water in a soil for any extended period is harmful
to plant growth because of the lack of air
Soil
Moisture Constants
Earlier classification divided soil water into
gravitational, capillary and hygroscopic water. The hygroscopic and capillary
waters are in equilibrium with the soil under given condition. The hygroscopic
coefficient and the maximum capillary capacity are the two equilibrium points
when the soil contains the maximum amount of hygroscopic and capillary waters,
respectively. The amount of water that a soil contains at each of these
equilibrium points is known as soil moisture constant.
The soil moisture constant, therefore, represents
definite soil moisture relationship and retention of soil moisture in the
field. The three classes of water (gravitational, capillary and hygroscopic)
are however very broad and do not represent accurately the soil - water relationships
that exists under field conditions.
- Though the maximum capillary capacity
represents the maximum amount of capillary water that a soil holds, the
whole of capillary water is not available for the use of the plants.
- A part of it, at its lower limit approaching
the hygroscopic coefficient is not utilized by the plants.
- Similarly a part of the capillary water at its
upper limit is also not available for the use of plants.
Hence two more soil constants, viz., field capacity
and wilting coefficient have been introduced to express the soil-plant-water
relationships as it is found to exist under field conditions.
1. Field capacity: Assume that water is applied to the surface of a soil. With
the downward movement of water all macro and micro pores are filled up. The
soil is said to be saturated with respect to water and is at maximum water
holding capacity or maximum retentive capacity. It is the amount of water held
in the soil when all pores are filled.
Sometimes, after application of water in the soil all
the gravitational water is drained away, and then the wet soil is almost
uniformly moist. The amount of water held by the soil at this stage is known as
the field capacity or normal moisture capacity of that soil. It is the capacity
of the soil to retain water against the downward pull of the force of gravity.
At this stage only micropores or capillary pores are filled with water and
plants absorb water for their use. At field capacity water is held with a force
of 1/3 atmosphere. Water at field capacity is readily available to plants and
microorganism
2. Wilting coefficient: As the moisture content
falls, a point is reached when the water is so firmly held by the soil
particles that plant roots are unable to draw it. The plant begins to wilt. At
this stage even if the plant is kept in a saturated atmosphere it does not
regain its turgidity and wilts unless water is applied to the soil. The stage
at which this occurs is termed the Wilting point and the percentage amount of
water held by the soil at this stage is known as the Wilting Coefficient. It
represents the point at which the soil is unable to supply water to the plant.
Water at wilting coefficient is held with a force of 15 atmospheres.
3. Hygroscopic coefficient: The hygroscopic
coefficient is the maximum amount of hygroscopic water absorbed by 100 g of dry
soil under standard conditions of humidity (50% relative humidity) and
temperature (15°C). This tension is equal to a force of 31 atmospheres. Water
at this tension is not available to plant but may be available to certain
bacteria.
4. Available water capacity: The amount of water
required to apply to a soil at the wilting point to reach the field capacity is
called the "available" water. The water supplying power of soils is
related to the amount of available water a soil can hold. The available water
is the difference in the amount of water at field capacity (- 0.3 bar) and the
amount of water at the permanent wilting point (- 15 bars).
5. Maximum water holding capacity: It is also known
as maximum retentive capacity. It is the amount of moisture in a soil when its
pore spaces both micro and macro capillary are completely filled with water. It
is a rough measure of total pore space of soil. Soil moisture tension is very
low between 1/100th to 1/1000th of an atmosphere or pF 1 to 0.
6. Sticky point moisture: It represents the
moisture content of soil at which it no longer sticks to a foreign object. The
sticky point represents the maximum moisture content at which a soil remains
friable. Sticky point moisture values vary nearly approximate to the moisture
equivalent of soils. Summary of the soil moisture constants, type of water and
force with which it held is given in following table.
Soil moisture constants and range of tension and PF
|
Sr. No.
|
Moisture
class
|
Range of
tension in atmosphere
|
Equivalent
PF range
|
|
1
|
Chemically
combined
|
Very High
|
-
|
|
2
|
Water vapour
|
Held at saturation point in the soil air
|
-
|
|
3
|
Hygroscopic
|
31 to 10,000
|
4.50 to 7.00
|
|
4
|
Hygroscopic
coefficient
|
31
|
4.50
|
|
5
|
Wilting point
|
15
|
4.20
|
|
6
|
Capillary
|
1/3 to 31
|
2.54 to 4.50
|
|
7
|
Moisture equivalent
|
1/3 to 1
|
2.70 to 3.00
|
|
8
|
Field capacity
|
1/3
|
2.54
|
Entry of
Water into Soil
1. Infiltration: Infiltration refers to the
downward entry or movement of water into the soil surface. It is a surface
characteristic and hence primarily influenced by the condition of the surface
soil. Soil surface with vegetative cover has more infiltration rate than bare
soil. Warm soils absorb more water than colder ones. Coarse surface texture,
granular structure and high organic matter content in surface soil, all help to
increase infiltration. Infiltration rate is comparatively lower in wet soils
than dry soils.
Factors affecting infiltration
1.
Clay minerals
2.
Soil Texture
3.
Soil structure
4.
Moisture content
5.
Vegetative cover
6.
Topography
2. Percolation: The movement of water through a
column of soil is called percolation. It is important for two reasons.
i) This is the only source of recharge of ground water which can be used
through wells for irrigation
ii) Percolating waters carry plant nutrients down and often out of reach of
plant roots (leaching)
Percolation
is dependent of rainfall. In dry region it is negligible and under high
rainfall it is high. Sandy soils have greater percolation than clayey soil.
Vegetation and high water table reduce the percolation loss
3. Permeability: It indicates the relative ease of
movement of water with in the soil. The characteristics that determine how fast
air and water move through the soil are known as permeability. The term
hydraulic conductivity is also used which refers to the readiness with which a
soil transmits fluids through it.
Soil Water Movement
i) Saturated Flow
ii) Unsaturated Flow
iii) Water Vapour Movement
Saturated flow: This occurs when the soil pores
are completely filled with water. This water moves at water potentials larger
than – 33 k Pa. Saturated flow is water flow caused by gravity’s pull. It
begins with infiltration,
which is water movement into soil when rain or irrigation water is on the soil
surface. When the soil profile is wetted, the movement of more water flowing
through the wetted soil is termed percolation.
Hydraulic conductivity can be expressed mathematically as
V
= kf
Where,
V = Total volume of
water moved per unit time
f = Water moving force
k = Hydraulic
conductivity of soil
Factors affecting movement of
water
1.
Texture, 2.Structure, 3.Amount of organic matter, 4.Depth of soil to hard pan,
5.Amount of water in the soil, 6.temperature and 7. Pressure
Vertical water flow:
The
vertical water flow rate through soil is given by Darcy’s law. The law
states that the rate of flow of liquid or flux through a porous medium is
proportional to the hydraulic gradient in the direction of floe of the liquid.
(dw) At
QW = - k--------------
Ds
Where,
QW = Quantity of water
in cm-3
k = rate constant (cm/s)
dw = Water height
(head), cm
A = Soil area (cm2))
t = Time
ds = Soil depth (cm)
Soil Air
Soil air is a continuation of the atmospheric air.
Unlike the other components, it is constant state of motion from the soil pores
into the atmosphere and from the atmosphere into the pore space. This constant
movement or circulation of air in the soil mass resulting in the renewal of its
component gases is known as soil aeration.
Composition of Soil Air: The soil air contains a
number of gases of which nitrogen, oxygen, carbon dioxide and water vapour are
the most important. Soil air constantly moves from the soil pores into the
atmosphere and from the atmosphere into the pore space. Soil air and atmospheric
air differ in the compositions. Soil air contains a much greater proportion of
carbon dioxide and a lesser amount of oxygen than atmospheric air. At the same
time, soil air contains a far great amount of water vapour than atmospheric
air. The amount of nitrogen in soil air is almost the same as in the
atmosphere.
Composition of soil and atmospheric air
|
Percentage by volume
|
|
|
Nitrogen
|
Oxygen
|
Carbon
dioxide
|
|
Soil Air
|
79.2
|
20.6
|
0.3
|
|
Atmospheric Air
|
79.9
|
20.97
|
0.03
|
Factors Affecting the Composition of Soil Air:
1. Nature and condition of soil: The quantity of
oxygen in soil air is less than that in atmospheric air. The amount of oxygen
also depends upon the soil depth. The oxygen content of the air in lower layer
is usually less than that of the surface soil. This is possibly due to more
readily diffusion of the oxygen from the atmosphere into the surface soil than
in the subsoil. Light texture soil or sandy soil contains much higher
percentage than heavy soil. The concentration of CO2 is usually greater in
subsoil probably due to more sluggish aeration in lower layer than in the
surface soil.
2. Type of crop: Plant roots require oxygen, which
they take from the soil air and deplete the concentration of oxygen in the soil
air. Soils on which crops are grown contain more CO2 than fallow lands. The
amount of CO2 is usually much greater near the roots of plants than further
away. It may be due to respiration by roots.
3. Microbial activity: The microorganisms in soil
require oxygen for respiration and they take it from the soil air and thus
deplete its concentration in the soil air. Decomposition of organic matter
produces CO2 because of increased microbial activity. Hence, soils rich in
organic matter contain higher percentage of CO2.
4. Seasonal variation: The quantity of oxygen is usually higher in dry
season than during the monsoon. Because soils are normally drier during the
summer months, opportunity for gaseous exchange is greater during this period.
This results in relatively high O2 and low CO2 levels. Temperature also
influences the CO2 content in the soil air. High temperature during summer
season encourages microorganism activity which results in higher production of
CO2.
Exchange of
Gases between Soil and Atmosphere
The exchange of gases between the soil and the
atmosphere is facilitated by two mechanisms
1. Mass flow: With every rain or irrigation, a part
of the soil air moves out into the atmosphere as it is displaced by the
incoming water. As and when moisture is lost by evaporation and transpiration,
the atmospheric air enters the soil pores. The variations in soil temperature
cause changes in the temperature of soil air. As the soil air gets heated
during the day, it expands and the expanded air moves out into the atmosphere.
On the other hand, when the soil begins to cool, the soil air contracts and the
atmospheric air is drawn in.
2. Diffusion: Most of the gaseous interchange in
soils occurs by diffusion. Atmospheric and soil air contains a number of gases
such as nitrogen, oxygen, carbon dioxide etc., each of which exerts its own
partial pressure in proportion to its concentration.
The movement of each gas is regulated by the
partial pressure under which it exists. If the partial pressure on one of the
gases (i.e. carbon dioxide) is greater in the soil air than in the atmospheric
air, it (CO2) moves out into the atmosphere. Hence, the concentration of CO2 is
more in soil air.
On the other hand, partial pressure of oxygen is
low in the soil air, as oxygen present in soil air is consumed as a result of
biological activities. The oxygen present in the atmospheric air (partial
pressure of O2 is greater) therefore, diffuses into the soil air till
equilibrium is established. Thus, diffusion allows extensive movement and
continual change of gases between the soil air and the atmospheric air. Oxygen
and carbon dioxide are the two important gases that take in diffusion
Importance
of Soil Aeration
1. Plant and root growth: Soil aeration is an
important factor in the normal growth of plants. The supply of oxygen to roots
in adequate quantities and the removal of CO2 from the soil atmosphere are very
essential for healthy plant growth.
When the supply of oxygen is inadequate, the plant
growth either retards or ceases completely as the accumulated CO2 hampers the
growth of plant roots. The abnormal effect of insufficient aeration on root
development is most noticeable on the root crops. Abnormally shaped roots of
these plants are common on the compact and poorly aerated soils. The
penetration and development of root are poor. Such undeveloped root system
cannot absorb sufficient moisture and nutrients from the soil
2. Microorganism population and activity: The
microorganisms living in the soil also require oxygen for respiration and
metabolism. Some of the important microbial activities such as the
decomposition of organic matter, nitrification, Sulphur oxidation etc depend
upon oxygen present in the soil air. The deficiency of air (oxygen) in soil
slows down the rate of microbial activity.
For example, the decomposition of organic matter is
retarded and nitrification arrested. The microorganism population is also
drastically affected by poor aeration.
3. Formation of toxic material: Poor aeration
results in the development of toxin and other injurious substances such as
ferrous oxide, H2S gas, CO2 gas etc in the soil.
4. Water and nutrient absorption: A deficiency of
oxygen has been found to check the nutrient and water absorption by plants. The
energy of respiration is utilized in absorption of water and nutrients. Under
poor aeration condition (this condition may arise when soil is water logged),
plants exhibit water and nutrient deficiency
5. Development of plant diseases: Insufficient
aeration of the soil also leads to the development of diseases. For example,
wilt of gram and dieback of citrus and peach.
Organic
Colloids – Soil Organic Matter
Soil
organic matter (SOM) comprises an accumulation of
i)
Partially disintegrated and decomposed plant and animal residues
ii) Other organic compounds synthesized by the soil microbes upon decay.
OM
content of a well drained mineral soil is LOW: 1 – 6 % by weight in the top
soil and even less in the subsoil.
Sources of Soil Organic Matter
The primary sources of SOM are plant tissues
1.
The tops and roots of trees
2.
Shrubs, grasses, remains of harvested crops and Soil organisms
Animals are secondary sources of Organic Matter.
1.
Waste products of animals
2.
Remains of animals after completion of life cycle.
Factors
Affecting Soil Organic Matter
1. Climate: Temperature and rainfall exert a
dominant influence on the amounts of N and OM found in soils.
a) Temperature: The OM and N content of comparable soils tend to increase if
one moves from warmer to cooler areas. The decomposition of OM is accelerated
in warm climates as compared to cooler climates. For each 10 C decline in mean
annual temperature, the total OM and N increases by two to three times.
b) Rainfall: There is an increase in OM with an increase in rainfall. Under
comparable conditions, the N and OM increase as the effective moisture becomes
greater.
2. Natural Vegetation: The total OM is higher in
soils developed under grasslands than those under forests.
3. Texture: Fine textured soils are generally
higher in OM than coarse textured soils.
4. Drainage: Poorly drained soils because of their
high moisture content and relatively poor aeration are much higher in OM and N
than well drained soils.
5. Cropping and Tillage: The cropped lands have
much low N and OM than comparable virgin soils. Modern conservation tillage
practices helps to maintain high OM levels as compared to conventional tillage.
6. Rotations, residues and plant nutrients: Crop
rotations of cereals with legumes results in higher soil OM. Higher OM levels,
preferably where a crop rotation is followed.
Decomposition
of Soil Organic Matter
The organic materials (plant and animal residues)
incorporated in the soil are attacked by a variety of microbes, worms and
insects in the soil if the soil is moist. Some of the constituents are
decomposed very rapidly, some less readily, and others very slowly. The list of
constituents in terms of ease of decomposition:
1. Sugars, starches and simple proteins Rapid Decomposition
2. Crude proteins
3. Hemicelluloses
4. Cellulose
5. Fats, waxes, resins
6. Lignins Very slow Decomposition
The organic matter is also classified on the basis
of their rate of decomposition
Rapidly decomposed: Sugars, starches, proteins etc
Less rapidly decomposed: Hemicelluloses, celluloses etc
Very slowly decomposed: Fats, waxes, resins, lignins etc
Simple decomposition products
Aerobic – CO2, H2O, NO3, SO4
When organic material is added to soil, three
general reactions take place
a. The bulk of the material undergoes enzymatic
oxidation with CO2, water, energy and heat as the major products.
b. The essential elements such as N, P and S are released and / or immobilized
by a series of reactions.
Compounds very resistant to microbial action is
formed either through modification of compounds or by microbial synthesis.
A. Decomposition of soluble substances: When
glucose is decomposed under aerobic conditions the reaction is as under:
Sugar + Oxygen --------------> CO2 + H2O
Under partially oxidized conditions,
Sugar + Oxygen --------------> Aliphatic acids
(Acetic, formic etc.) or
Hydroxy acids
(Citric, lactic etc.) or
Alcohols (ethyl alcohol etc.)
Some of the reactions invoiced may be represented
as under:
C6H12O6 + 2O2 -------------> 2 CH3. COOH + 2CO2 + 2H2O
2C6H12O6 + 3O2 ------------> 2 C6H8O7 + 4 H2O
C6H12O6 + 2O2 -------------> 2C2H5OH + 2 CO2
i) Ammonification: The transformation of organic
nitrogenous compounds (amino acids, amides, ammonium compounds, nitrates etc.)
into ammonia is called ammonification. This process occurs as a result of
hydrolytic and oxidative enzymatic reaction under aerobic conditions by
heterotrophic microbes.
ii)Nitrification: The process of conversion of
ammonia to nitrites (NO2) and then to nitrate (NO-3) is known as nitrification.
It is an aerobic process by autotrophic bacteria.
Nitrosomonas
Nitrobacter
NH3 --------------------------> NO2 -------------------> NO-3
Ammonia
Nitrite
Nitrate
The net reactions are as follows:
NH4 + O2 -------------------> NO2 + 2H+ + H2O + energy
NO2 + O2 ------------------- >NO-3 + energy
iii) Denitrification: The process, which involves
conversion of soil nitrate into gaseous nitrogen or nitrous oxide, is called
Denitrification. Water logging and high pH will increase N loss by
Denitrification.
Pseudomonas / Bacillus
Nitrate --------------------------------> Nitrogen Gas
B. Decomposition of Insoluble Substances
i) Breakdown of Protein: During the course of
decomposition of plant materials, the proteins are first hydrolyzed to a number
of intermediate products and may be represented as under:
Hydrolysis
Proteases Aas
Proteins --------------> Peptones --------> Amides ------> Ammonia
Peptides
Aminization
Ammonification
Aminization: The process of conversion of proteins to aminoacids.
Ammonification: The process of conversion of aminoacids and amides to ammonia.
ii) Breakdown of cellulose: The decomposition of
the most abundant carbohydrates is as follows:
hydrolysis
hydrolysis
Cellulose ---------------> Cellobiose ----------------> Glucose
(cellulase)
(cellobiase)
oxidation
---------------> Organic acids -------------> CO2 + H2O
This reaction proceeds more slowly in acid soils than in neutral and alkaline
soils. It is quite rapid in well aerated soils and comparatively slow in poorly
aerated soils.
iii) Breakdown of Hemicellulose: Decompose faster
than cellulose and are first hydrolyzed to their components sugars and uronic
acids. Sugars are attacked by microbes and are converted to organic acids,
alcohols, carbon dioxide and water. The uronic acids are broken down to pentose
and CO2. The newly synthesized hemicelluloses thus form a part of the humus.
iv) Breakdown of Starch: It is chemically a glucose
polymer and is first hydrolyzed to maltose by the action of amylases. Maltose
is next converted to glucose by maltase. The process is represented as under:
(C6H10O5)n +nH2O -------------> (C6H12O6)
C. Decomposition of ether soluble substances:
Fats------------------------> glycerol + fatty
acids
Glycerol -------------------> CO2 + water
D. Decomposition of lignin: Lignin decomposes slowly, much slower than
cellulose. Complete oxidation gives rise to CO2 and H2O.
Factors
affecting decomposition of organic matter
1. Temperature: Cold periods retard plant growth
and OM decomposition. Warm summers may permit plant growth and humus
accumulation.
2. Soil moisture: Extremes of both arid and
anaerobic conditions reduce plant growth and microbial decomposition. Near or
slightly wetter than field capacity moisture conditions are most favorable for
both processes.
3. Nutrients: Lack of nutrients particularly N
slows decomposition.
4. Soil PH: Most of the microbes grow best at pH 6
– 8, but are severely inhibited below pH 4.5 and above pH 8.5.
5. Soil Texture: Soils higher in clays tend to
retain larger amounts of humus.
6. Other Factors: Toxic levels of elements (Al, Mn,
B, Se, Cl), excessive soluble salts, shade and organic phytotoxins in plant
materials.
Role of
organic matter
1. Organic Matter creates a granular condition of
soil which maintains favorable condition of aeration and permeability.
2. Water holding capacity of soil is increased and
surface runoff, erosion etc., are reduced as there is good infiltration due to
the addition of OM.
3. Surface mulching with coarse OM lowers wind
erosion and lowers soil temperatures in the summer and keeps the soil warmer in
winter.
4. OM serves as a source of energy for the microbes
and as a reservoir of nutrients that are essential for plant growth and also
hormones, antibiotics.
5. Fresh OM supplies food for earthworms, ants and
rodents and makes soil P readily available in acid soils.
6. Organic acids released from decomposing OM help
to reduce alkalinity in soils; organic acids along with released CO2 dissolve
minerals and make them more available.
7. Humus (a highly decomposed OM) provides a
storehouse for the exchangeable and available cations.
8. It acts as a buffering agent which checks rapid
chemical changes in pH and soil reaction.
Humus
Humus is a complex and rather resistant mixture of
brown or dark brown amorphous and colloidal organic substance that results from
microbial decomposition and synthesis and has chemical and physical properties
of great significance to soils and plants.
Humus Formation
The humus compounds have resulted from two general
types of biochemical reactions: Decomposition and Synthesis.
1. Decomposition:
a) Chemicals in the plant residues are broken down by soil microbes including
lignin.
b) Other simpler organic compounds that result from the breakdown take part
immediately in the second of the humus-forming processes, biochemical
synthesis.
c) These simpler chemicals are metabolized into new compounds in the body
tissue of soil microbes.
d) The new compounds are subject to further modification and synthesis as the
microbial tissue is subsequently attacked by other soil microbes.
2. Synthesis: Involve such breakdown products of
lignin as the phenols and quinones.
a) These monomers undergo polymerization by which polyphenols and polyquinones
are formed.
b) These high molecular weight compounds interact with N-containing amino
compounds and forms a significant component of resistant humus.
c) Colloidal clays encourage formation of these polymers.
d) Generally two groups of compounds that collectively make up humus, the humic
group and the nonhumic group.
Properties
of Humus
1. The tiny colloidal particles are composed of C,
H, and O2. The colloidal particles are negatively charged (-OH, -COOH or
phenolic groups), has very high surface area, higher CEC (150 – 300 cmol/kg), 4
- 5 times higher WHC than that of silicate clays.
2. Humus has a very favorable effect on aggregate
formation and stability.
3. Impart black colour to soils.
4. Cation exchange reactions are similar to those
occurring with silicate clays.
Clay – Humus Complex
Humus, the organic amorphous colloid supplies both
basic and acidic ions which is transitory and ultimately disappears from soil.
Clay, the inorganic crystalline colloid supplies chiefly the basic nutrient
ions is more or less stable. Both these colloids form the soil colloidal
complex and are extremely active and form important sources of plant nutrients.
It is believed that humus and clay exist in the soil as clay – humus complex,
the two being held together by cations like Ca, Fe, etc. Depending upon the
nature of binding cation, two types of Clay – humus complex have been
recognized. The colloidal complex bound by Ca ions is more stable and is
responsible for the favorable physical condition of the soil, particularly its
structure. The other type where Fe acts as the binding agent creates a poor
physical condition of the soils.
Ion
Exchange
As soils are formed during the weathering
processes, some minerals and organic matter are broken down to extremely small
particles. Chemical changes further reduce these particles until they cannot be
seen with the naked eye. The very smallest particles are called colloids.
The mineral clay colloids are plate like in
structure and crystalline in nature. In most soils; clay colloids exceed
organic colloids in amount.
Colloids are primarily responsible for the chemical
reactivity in soils. The kind of parent material and the degree of weathering
determine the kinds of clays present in the soil. Since soil colloids are
derived from these clays, their reactivity is also influenced by parent
material and weathering. Each colloid (inorganic and organic) has a net
negative (-) charge developed during the formation process. This means it can
attract and hold positively (+) charged particles. An element with an
electrical charge is called an ion.
Potassium, sodium (Na), hydrogen (H), Ca and Mg all
has positive charges. They are called cations and ions with negative charges,
such as nitrate and sulfate, are called anions.
Negatively charged colloids attract cations and
hold them like a magnet holds small pieces of metal. This characteristic
explains; why nitrate-N is more easily leached from the soil than ammonium-N.
Nitrate has a negative charge like soil colloids. So, NO3- is not held by the
soils, but remains as a free ion in soil water to be leached through the soil
profile in some soils and under some rainfall conditions. The charges
associated with soil particles attract simple and complex ions of opposite
charge. Thus, a given colloidal mixture may exhibit not only a maze of positive
and negative surface charges but also an equal complex complement of simple
cations and anions such as Ca2+ and S04 - that are attracted by the particle
charges.
The adsorbed anions are commonly present in smaller
quantities than the cations because the negative charges generally predominate
on the soil colloid.
Mechanism of Cation Exchange: The exchange of
cations has been explained on the basis of the electro-kinetic theory of ion
exchange. According to this theory, the adsorbed cations forming the outer
shell of the ionic double layer are supposed to be in a state of oscillation
when suspended in water, forming a diffuse double layer. Due to these
oscillations, some of the cations move away from the surface of the clay
micelle. In the presence of the solution of an electrolyte a cation of the
added electrolyte slips in between the inner negative layer and the outer
oscillating positive ion. The electrolyte cation is now adsorbed on the micelle
and the surface cation remains in solution as an exchanged ion. Thus the
exchange of cations takes place. Cations are positively charged nutrient ions
and molecules. While, clay particles are negatively charged constituents of
soils. These negatively charged particles (clay) attract, hold and release
positively charged nutrient ions (cations). Organic matter particles also have
a negative charge to attract cations. Sand particles carry little or no charge
and do not react. Cations held by soils can be replaced by other cations. This
means they are exchangeable. For example, Ca++ can be exchanged for H+ and /or
K+ and vice versa.
Cation Exchange Capacity (CEC): The CEC is the capacity of soil to hold and
exchange cations. The cation exchange capacity is defined simply as the sum
total of the exchangeable cations that a soil can adsorb. The higher the CEC of
soil the more cations it can retain. Soils differ in their capacities to hold
exchangeable K+ and other cations.
The CEC depends on amount and kinds of clay and
organic matter present. A high-clay soil can hold more exchangeable cations
than a low-clay soil. CEC also increases as organic matter increases. Clay
minerals usually range from 10 to 150 meq/100 g in CEC values. Organic matter
ranges from 200 to 400 meq/100 g. So, the kind and amount of clay and organic
matter content greatly influence the CEC of soils. Clay soils with high CEC can
retain large amounts of cations against potential loss by leaching. Sandy
soils, with low CEC, retain smaller quantities.
This makes timing and application rates important
in planning a fertilizer programme. For example, it may not be wise to apply K
on very sandy soils in the middle of a monsoon, where rainfall can be high and
intense. Fertilizer application should be split to prevent leaching and losses
through erosion. Also, splitting N applications to meet peak crop demand are
important to reduce the potential for nitrate leaching on sands as well as
finer-textured soils.
Means of Expression: The cation exchange capacity
is expressed in terms of equivalents or more specifically, as milli equivalents
per 100 gram and is written as meq /100g. The term equivalent is defined as one
gram of atomic weight of hydrogen (or the amount of any other ion) that will
combine with or displace this amount of hydrogen for monovalent ions such as
Na+, K+, NH4+ and Cl -, the equivalent weight and atomic weight are same, since
they can replace one H ion. Divalent cations such as Ca++ and Mg++ can take the
place of two H+ ions. The milliequivalent weight of a substance is one
thousandth of its atomic weight. Since the atomic weight of hydrogen is about 1
gram. The term milliequivalent (meq.) may be defined as 1 milligram of
hydrogen.
This unit of exchangeable cations i.e. milliequivalent per 100 g of soil
(meq/100g) was used prior to 1982. In the newer metric system the term
equivalent is not used, however, now moles are the accepted chemical unit. All
the calculation and concepts of "equivalents" are still mentally used
but the notation must be written differently. The old "equivalent" is
represented by moles (+) or mole, which indicates a monovalent ion portion. For
example, to write 12.5 meq/100 g in the newer metric system, it can be written
as: 12.5 c mol (+) kg-1 of soil (centimoles) or 125 m mol (+) kg-1 of soil
(millimoles).
Replacing power of cations: The replacing power of
cations varies with the type of ion, its size and degree of hydration, valence
and concentration and the kind of clay mineral involved, as it is controlled by
number of factors no single order of replacement can be given. All other
factors being equal the replacing power of monovalent cations increases in the
following order: Li < Na < K < Rb < Cs < H and for divalent
cations: Mg < Ca < Sr < Ba. In case of mixture of monovalent and
divalent cations as they exist in normal soils the replacing power increases in
the following order: Na < K < NH4 < Mg < Ca < H. This means Na
is more easily replaced than K and K more easily than NH4 and so on.
Percent base saturation: The percent of total CEC
occupied by the major cations has been used in the past to develop fertilizer
programs. The idea is that certain nutrient ratios or 'balances' are needed to
ensure proper uptake by the crop for optimum yields. Research has shown,
however, that cation saturation ranges and ratios have little or no utility in
a vast majority of agricultural soils. Under field conditions, ranges of
nutrients can vary widely with no detrimental effects, so long as individual
nutrients are present in sufficient levels in the soil to support optimum plant
growth.
Importance of Cation Exchange: Cation exchange is
an important reaction in soil fertility, in causing and correcting soil acidity
and basicity, in changes altering soil physical properties, and as a mechanism
in purifying or altering percolating waters. The plant nutrients like calcium,
magnesium, and potassium are supplied to plants in large measure from
exchangeable forms.
Soil
reaction
Soil reaction is one of the most important
physiological characteristics of the soil solution. The presence and
development of micro- organisms and higher plants depend upon the chemical
environment of soil. There fore study of soil reaction is important in soil
science.
There are three types of soil reactions: 1. Acidic
2. Alkaline and 3. Neutral
1. Acidic: It is common in region where
precipitation is high. The high precipitation leaches appreciable amounts of
exchangeable bases from the surface layers of the soils so that the exchange
complex is dominated by H ions. Acid soils, therefore, occur widely in humid
regions and affect the growth of plants markedly.
2. Alkaline: Alkali soils occur when there is
comparatively high degree of base saturation. Salts like carbonates of calcium,
magnesium and sodium also give a preponderance of OH ions over H ions in the
soil solution. When salts of strong base such as sodium carbonate go into soil
solution and hydrolyze, consequently they give rise to alkalinity. The reaction
is as follows:
Na2CO3 -----à 2Na + + CO3=
2Na+ + CO3= + 2HOH -----à 2Na+ + 2OH - + H2CO3
since sodium hydroxide dissociates to a greater degree than the carbonic acid,
OH ions dominate and give rise to alkalinity. This may be as high as 9 or 10.
These soils most commonly occur in arid and semi-arid regions.
3. Neutral: Neutral soils occur in regions where H
ions just balance OH ions.
Soil pH: The reaction of a solution represents the
degree of acidity or basicity caused by the relative concentration of H ions
(acidity) or OH ions present in it. Acidity is due to the excess of H ions over
OH ions, and alkalinity is due to the excess of OH ions over H ions. A neutral
reaction is produced by an equal activity of H and OH ions. According to the
theory of dissociation, the activity is due to the dissociation or ionization
of compounds into ions.
Method of
expressing acidity or alkalinity
Equivalent quantities of all acids or alkalies
contain the same number of total H or OH ions, respectively. But, when they are
dissolved in water they do not ionize to the same extent. The amount of acid or
alkali ionized depends upon the content of free H and OH ions. When the
dissociation is high e.g., hydrochloric acid (a strong acid), it dissociates to
a larger extent than the weak acetic acid. Acetic acid dissociates only to
about 10 % as compared to hydrochloric acid. In a 1N solution of hydrochloric
acid there will be 1 gram of H ions per liter, while in a normal solution of
acetic acid there will be 1/10 gram of H ions per liter. But in titration 1 ml
of 1 N hydrochloric acid and 1 ml of 1 N acetic acid will require 1 ml of 1 N
sodium hydroxide for neutralization separately, because the total acidity is
the same and titration determines both the ionized and unionized H or OH ions.
Soil acidity: There are three kinds of acidity.
(i) Active acidity is due to the H+ ion in the soil solution.
(ii) Salt replaceable acidity represented by the hydrogen and aluminum that are
easily exchangeable by other cations in a simple unbuffered salt solution such
as KCl and
(iii) Residual acidity is that acidity which can be neutralized by limestone or
other alkaline materials but cannot be detected by the salt-replaceable
technique. Obviously, these types of acidity all add up to the total acidity of
a soil.
i. Active acidity: The active acidity is a measure
of the H+ ion activity in the soil solution at any given time. However, the
quantity of H+ ions owing to active acidity is very small compared to the
quantity in the exchange and residual acidity forms. For example, only about 2
kg of calcium carbonate would be required to neutralize the active acidity in a
hectare-furrow slice of an average mineral soil at pH 4 and 200/0 moisture.
Even though the concentration of hydrogen ions owing to active acidity is
extremely small, it is important because this is the environment to which
plants and microbes are exposed.
ii. Salt replaceable (exchangeable) acidity: This
type of acidity is primarily associated with the exchangeable aluminum and
hydrogen ions that are present in largest quantities in very acid soils. These
ions can be released into the soil solution by an unbuffered salt such as KCl.
Al3+ + 4KCI +AlCI3+HCI' L~~~ H + L~~~~ ~ (Soil Solid) (Soil Solution) (Soil
Solid) (Soil solution), in moderately acid soils, the quantity of easily
exchangeable aluminum and hydrogen is quite limited. Even in these soils,
however, the limestone needed to neutralize this type of acidity is commonly
more than 100 times that needed for the soil solution (active acidity). At a
given pH value, exchangeable acidity is generally highest for smectites,
intermediate for vermiculites, and lowest for kaolinite. In any case, however,
it accounts for only a small portion of the total soil acidity as the next
section will verify.
iii. Residual acidity: Residual acidity is that
which remains in the soil after active and exchange acidity has been
neutralized. Residual acidity is generally associated with aluminum hydroxy
ions and with hydrogen and aluminum atoms that are bound in non exchangeable
forms by organic matter and silicate clays. If lime is added to a soil, the pH
increases and the aluminum hydroxy ions are changed to uncharged gibbsite as
follows. OH- OH- AI (OH) 2+ 7 AI (OH) 2+ 7 AI (OH) 3 In addition, as the pH
increases bound hydrogen and aluminum can be released by calcium and magnesium
In the lime materials [Ca (OH) 2 is used as an example of the reactive calcium
liming material] The residual acidity is commonly far greater than either the
active or salt replaceable acidity. Conservative estimates suggest that the
residual acidity may be 1000 times greater than the soil solution or active
acidity in a sandy soil and 50000 or even 100000 times greater in a clayey soil
high in organic matter. The amount of ground limestone recommended to at least
partly neutralize residual acidity is commonly 4-8 metric tons (Mg) per hectare
furrow slice (1.8-3.6 tons/AFS).
It is obvious that the pH of the soil solution is only "the tip of the
iceberg" in determining how much lime is needed. Buffering and Soil
Reaction Buffer action: Buffering refers to resistance to a change in pH. If 1
ml HCI (of 0.1 N) is added to one liter of pure distilled water of pH 7.0, the
resulting solution would have a pH of about 5.0. If on the other hand, the same
amount of acid is added to a liter of soil suspension the resulting change in
pH would be very small. There is, a distinct resistance to change in pH. This
power to resist a change in pH is called buffer action. A buffer solution is
one which contains reserve acidity and alkalinity and does not change pH with
small additions of acids or alkalies. Buffer capacity: The colloidal complex
acts as a powerful buffer in the soil and does not allow rapid and sudden
changes in soil reaction. Buffering depends upon the amount of colloidal
material present in soil. Clay soils rich in organic matter are more highly
buffered than sandy soils. Buffer capacity of the soil varies with its cation
exchange capacity (C.E.C.). The greater the C.E.C. the greater will be its
buffer capacity. Thus heavier the texture and the greater the organic matter
content of a soil, the greater is the amount of acid or alkaline material
required to change its pH. Importance of Buffering to Agriculture: Changes In
soil reaction (pH) have a direct influence on the plants and it also affects
the availability of plant nutrients. Deficiency of certain plant nutrients and
excess availability of others in toxic amounts would seriously upset the
nutritional balance in the soil. Buffering prevents sudden changes and
fluctuation in soil pH, so it regulates the availability of nutrients and also
checks direct toxic effect to plants.
Factors
Controlling Soil Reactions
Soil reaction varies due to following factors
1. Nature of soil colloids: The colloidal particles
of the soil influence soil reaction to a very greatest extent. When hydrogen
(H+) ion forms the predominant adsorbed cations on clay colloids, the soil
reaction becomes acid.
2. Soil solution: The soil solution carries a
number of salts dissolved in capillary water. The cations of the salts
intermingle with those of the diffuse double layer of the clay particle and
increase the concentration. The concentration of cations in bulk of the
solution is more or less (or nearly) the same as that near the particle
surfaces. For an unsaturated soil (Clay), the more compact the layer the
greater is the number of hydrogen ions dissociating into the solution. This
increases the acidity of the soil solution or lowers its pH. Under field
conditions, the concentration of salts varies with the moisture content of the
soil. The more dilute the solution, the higher the pH value. Hence the pH tends
to drop as the soil gets progressively dry. Soil reaction is also influenced by
the presence of CO2 in soil air. As the CO2 concentration increases, the soil
pH falls and increases the availability of the nutrients. Under field
conditions, plant roots and micro-organism liberate enough CO2 , which results
in lowering the pH appreciably. This principle of increasing the concentration
of CO2 in soil air is also used in the reclamation of alkali soils.
3. Climate: Rainfall plays important role in determining
the reaction of soil. In general, soils formed in regions of high rainfall are
acidic (low pH value), while those formed in regions of low rainfall are
alkaline (high pH value).
4. Soil management: Cultural operations in general
tend to increase soil acidity. They make an acid soil more acidic, and an
alkaline soil less alkaline. As a result of constant cultivation, basic
elements are lost from the soil through leaching and crop removal. This leads
to change the soil reaction to the acid side.
5. Parent materials: Soils developed from parent
material of basic rocks generally have higher pH than those formed from acid
rocks (e.g. granite). The influence of parent material is not very important as
it is completely masked by the climatic conditions under which the soil is
developed.
6. Precipitation: As water from rainfall passes through the soil, basic
nutrients such as calcium (Ca) and magnesium (Mg) are leached. They are
replaced by acidic elements including Al, H and manganese (Mn). Therefore, soils
formed under high rainfall conditions are more acid than those formed under
arid conditions.
7. Decomposition of
organic matter: Soil organic
matter is continuously being decomposed by micro-organisms into organic acids,
carbon dioxide (CO2) and water, forming carbonic acid. Carbonic acid, in turn,
reacts with the Ca and Mg carbonates in the soil to form more soluble
bicarbonates, which are leached away, leaving the soil more acid.
8. Native
vegetation: Soils often become
more acid when crops are harvested because of removal of bases. Type of crop
determines the relative amounts of removal. For example, legumes generally
contain higher levels of bases than do grasses. Calcium and Mg contents also
vary according to the portion (s) of the plant harvested. Many legumes release
H ions into their rhizosphere when actively fixing atmospheric N2. The acidity
generated can vary from 0.2 to 0.7 pH units per mole of fixed N.
9. Soil depth: Except in low rainfall areas, acidity generally
increases with depth, so the loss of topsoil by erosion can lead to a more acid
pH in the plough layer. The reason is that more subsoil is included in the plow
layer as topsoil is lost. There are areas, however, where subsoil pH is higher
than that of the topsoil.
10. Nitrogen fertilization: Nitrogen from fertilizer, organic matter, and
manure and legume N fixation produces acidity. Nitrogen fertilization speeds up
the rate at which acidity develops. At lower N rates, acidification rate is
slow, but is accelerated as N fertilizer rates increase.
11. Flooding: The overall effect of submergence is an increase
of pH in acid soils and a decrease in basic soils. Regardless of their original
pH values, most soils reach pH of 6.5 to 7.2 within one month after flooding
and remain at the level until dried. Consequently, liming is of little value in
flooded rice production. Further, it can induce deficiencies of micronutrients
such as zinc (Zn).
Influence
of Soil Reaction on Availability of Nutrients
The unproductiveness of acid and alkali soils is
very often due to the lack of available plant nutrients. In highly acid soils
(low pH), the availability of some of the nutrients such as aluminum, iron,
manganese etc., is increased to a point to become toxic to the plant. At the
same time the supplies of available calcium, nitrogen, phosphorus etc., are
reduced to starvation level (become unavailable). The same is the case at high
pH (alkaline conditions), plant growth suffers due to the unavailability of
nutrients like nitrogen, phosphorus and some minor elements (e.g., iron,
manganese, boron etc). Another indirect effect occurs through the activity of
microorganisms. Most microorganisms function at their best within a pH range
6.0 to 7.5. If soil reaction is changed beyond this range, the microorganisms
become functionless. Consequently the supply of some of the essential plant
nutrients like nitrogen is considerably reduced.
1. Nitrogen: Plant absorbs most of their nitrogen
in the form of nitrate of which availability depends on the activity of
nitrifying bacteria. The micro- organisms responsible for nitrification are
most active when the pH is between 6.5 and 7.5. They are adversely affected if
the pH falls below 5.5 and rises above 9.0. Nitrogen fixing bacteria (like
Azotobactor) also fail to function below pH 6.0. The decomposition of organic
matter which is the primary source of nitrogen is also slowed down under acidic
condition.
2. Phosphorus: Its availability is at its highest
when the reaction is between 6.5 and 7.5. When the reaction is above or below
this range, availability is reduced. In the strongly acidic soil (pH 5.0 or
less), iron, aluminum, manganese and other bases are present in a soluble state
and in more quantity. The phosphates of these elements are formed and become
unavailable.
3. Potassium: The availability of potassium does
not influence by soil reaction to any great extent. In acid soil potassium is
lost through leaching. The unavailability of K is due to the conversion of
exchangeable to non-exchangeable potassium in alkaline soil. Particularly if
the alkalinity is due to CaCO3 (brought about by over liming in acid soil), the
solubility of soil potassium is depressed.
4. Calcium and magnesium: Acid soils (base
unsaturated) are poor in / available calcium and magnesium. In alkaline soil
(pH not exceeding 8.5) the availability of Ca and Mg nutrients are always high.
When the pH is above 8.5, the availability of these nutrients again decreases.
5 Iron, aluminum and manganese: When the pH is low the solubility of
"iron, aluminum and manganese compounds are increased. and hence they are
readily available in acid soils. At the pH range 5.5 to 7.0, iron and manganese
are present in the soluble ferrous (Fe++) and manganous (Mn++) forms. At pH
below 5.5 the solubility of these compounds considerably increased with the
result that they have a toxic influence on plant growth. Under neutral and
alkaline conditions, iron and manganese are usually present in ferric (Fe3+ )
and manganese (Mn++++) states. Hence in soils with pH 7.5 and above, they
become unavailable and sometimes produce deficiency diseases like chlorosis in
plants.
6. Sulphur: The availability of sulphur is not
affected by soil reaction as sulphur compounds are soluble. in low pH range.
However, it is more soluble in acid soil and lost in leaching. Acid conditions,
which retard the decomposition of organic matter, therefore, retard the release
of available sulphur. The availability of sulphur present in organic matter
depends upon the decomposition of organic matter.
7. Micronutrients: In general, the availability of
boron, copper and zinc is reduced in alkaline soils and that of molybdenum in
acid soils. The availability of these nutrients progressively decreases as the
soil pH increases. Their availability also decreases under highly acid condition
when the pH is below 5.0. Zinc availability in alkaline soils from insoluble
zinc salts (calcium zincates) is reduced. Zinc and copper are adsorbed on the
clay colloids and not easily displaced and hence not available for plant
growth. The availability of molybdenum is reduced under acid soils. It is more
available in neutral and alkaline soils.
