SOIL SCIENCE LECTURE NOTES

(2014)

By

Jacqueline Abbo

TABLE OF CONTENTS

1. INTRODUCTION....................................................................................................3

1.1 Composition of soil..............................................................................................31.1.1 Soil Minerals........................................................................................................................................................31.1.2 Soil organic matter..............................................................................................................................................4

1.2 Importance of Soils.............................................................................................42. SOIL FORMATION AND DEVELOPMENT...................................................................5

2.1 Soil Formation Factors........................................................................................52.1.1 Soil Parent Material.............................................................................................................................................52.1.2 Climate................................................................................................................................................................52.1.3 Topography (Relief).............................................................................................................................................62.1.4 Biotic Activity......................................................................................................................................................62.1.5 Time....................................................................................................................................................................6

2.2 Formation of Inorganic Soils................................................................................72.2.1 Rocks..................................................................................................................................................................72.2.2 Physical weathering............................................................................................................................................72.2.3 Chemical weathering..........................................................................................................................................82.2.4 Inter-relationship between Physical and Chemical Weathering...........................................................................82.2.5 Rate of weathering..............................................................................................................................................9

3. THE SOIL PROFILE...............................................................................................11

3.1 Master horizons and layers................................................................................114. BASIC PHYSICAL PROPERTIES..............................................................................12

4.1 Soil Texture......................................................................................................124.1.1 Soil textural classes..........................................................................................................................................124.1.2 Determination of soil texture............................................................................................................................14

4.2 Soil Structure....................................................................................................154.2.1 Spherical Peds...................................................................................................................................................164.2.2 Rectilinear Peds................................................................................................................................................164.2.3 The Structure Grade..........................................................................................................................................174.2.4 Soil structure formation....................................................................................................................................184.2.5 Soil structure stability.......................................................................................................................................18

4.3 Soil Colour........................................................................................................184.3.1 Colour as a Guide to Soil Use............................................................................................................................194.3.2 Soil colour description.......................................................................................................................................194.3.3 The Munsell system of colour notation..............................................................................................................194.3.4 Munsell Soil Color Charts...................................................................................................................................20

4.4 Soil Density, ρ...................................................................................................214.4.1 Particle density ρs.............................................................................................................................................214.4.2 Bulk density, ρb (dry density)............................................................................................................................214.4.3 Soil Porosity, p..................................................................................................................................................214.4.4 Permeability......................................................................................................................................................224.4.5 Total (wet density), ρt.......................................................................................................................................224.4.6 Void ratio, e.......................................................................................................................................................224.4.7 Soil wetness, θ..................................................................................................................................................22

4.5 Stages of Soil Formation....................................................................................235. CHEMICAL PROPERTIES OF SOIL..........................................................................24

5.1 Soil pH:.............................................................................................................245.1.1 Measuring Soil pH.............................................................................................................................................245.1.2 How pH Affects Nutrients, Minerals and Growth................................................................................................245.1.3 Changes in Soil pH............................................................................................................................................26

5.2 Salinity (EC)......................................................................................................265.2.1 How soil salinity affects plant growth................................................................................................................265.2.2 Factors affecting soil salinity.............................................................................................................................275.2.3 Prevention and Management of Soil Salinity.....................................................................................................27

5.3 Cation exchange capacity (CEC).........................................................................295.3.1 Base Saturation.................................................................................................................................................305.3.2 CEC and Availability of Nutrients.......................................................................................................................305.3.3 Anion Exchange................................................................................................................................................30

5.4 Organic matter..................................................................................................315.4.1 Non- humified & humified organic matter:........................................................................................................315.4.2 Sources of organic matter.................................................................................................................................31

5.5 Plant tissue.......................................................................................................315.6 Animals.............................................................................................................31

pg. 1

5.7 Soil organisms..................................................................................................315.7.1 Types of materials present in organic matter:..................................................................................................325.7.2 Composition of plant tissues:............................................................................................................................32

5.8 C:N ratio (Carbon to Nitrogen)...........................................................................335.8.1 Importance of the C:N ration in soil...................................................................................................................345.8.2 Calculation Procedures of Carbon Nitrogen Ratio..............................................................................................34

6. THE COLLOIDAL NATURE OF SOIL........................................................................36

6.1 Soil Colloids......................................................................................................366.1.1 Classification of colloids....................................................................................................................................36

6.2 Clays................................................................................................................366.2.1 Mineralogy of clay.............................................................................................................................................366.2.2 Groups of silicate clays.....................................................................................................................................386.2.3 Sources of negative charge on silicate clay mineral.........................................................................................396.2.4 Non-Crystalline Silicate Clays............................................................................................................................406.2.5 Oxide Clays.......................................................................................................................................................406.2.6 Cation Exchange...............................................................................................................................................40

pg. 2

1. INTRODUCTIONThe definition of soil varies depending on the person considering it. To a civil engineer planning a construction site, soil is whatever unconsolidated material happens to be found at the surface. To a miner, it is just some worthless material that is in the way and must be removed in order to access the required mineral. To a farmer, it is the medium that will nourish and supply water to the crops. Even soil scientists may hold differing definitions, depending on their area of study.

The knowledge of soil science is aimed at introducing concepts and principles so that afterwards, one is able to appreciate soils as media for plant growth, reactor in waste disposal and foundation for roads and buildings, to mention but a few. There are different types of soil study, namely pedlogical study (The study of soil in relation to its characteristics, formation and distribution without emphasizing its utility or uses) and edaphological study (The study of soil emphasizing the organisms interactions with man as a major manipulating player). In this course, we shall mostly focus on the former and less on the latter.

Soil comprises of particles up to about 2mm diameter. They have horizons, roughly parallel to the earth’s surface, that indicate the degree to which materials have been altered and redistributed by water, gravity, and organisms. Individual soils occupy distinct places in the landscape, so we find patterns of soils and landscape features.

Landscapes and soils on them, together with the biota, are arranged and respond to their temporal and spatial distributions of water, nutrients and energy. And in particular, landscapes control the distribution of energy, water and nutrients.

Soil properties reflect the progressive alteration and redistributions of nutrients, minerals and organic matter over time.

1.1 Composition of soilSoil, in all stages of formation right from the parent material, is made up of two types of materials usually mixed together, namely organic materials and inorganic materials.

Soil comprises of soil-air/atmosphere, the soil-water (hydrosphere1), mineral-material (lithosphere2) and

organic-material or organisms (Biosphere3). These constituents are often referred to as soil phases. The soil phases are dynamic, i.e. they change with time and space and they are interactive. The ratio of air-filled pore space to water-filled pore space often changes seasonally, weekly, and even daily, depending on water additions through precipitation, throughflow, groundwater discharge, and flooding. The volume of the pore space itself can be altered, one way or the other, by several processes.

1.1.1 Soil MineralsAlmost any mineral that exists may be found in some soil, somewhere. The broad and deep subject area of soil mineralogy shall not be discussed here in detail; Only some of the elementary basics shall be discussed.

The mineral portion of soil is divided into three particle-size classes: sand, silt, and clay. [Note: Sand, silt, and clay are collectively referred to as the fine earth fraction of soil. They are <2 mm in diameter. Larger soil particles are referred to as rock fragments and have their own size classes (pebbles, cobbles, and boulders).

Mineralogically, sand, and silt are just small particles of rock and are largely inert. The two important differences among them are their relative capacity to hold water that is available for uptake by plants and their effects on soil drainage.

1Hydrosphere is the portion of Earth's surface that is water, including the seas and water in the atmosphere 2 Lithosphere is the solid outer layer of the Earth above the asthenosphere, consisting of the crust and upper mantle3 Biosphere is the whole area of Earth's surface, atmosphere, and sea that is inhabited by living things

pg. 3

Clay particles are mineralogically different from sand and silt. We shall discuss these details later.

1.1.2 Soil organic matterSoil organic matter (SOM) is a complex mixture of substances that can be highly variable in its chemical content. It ranges from freshly deposited plant and animal parts to the residual humus —stable organic compounds that are relatively resistant to further rapid decomposition.

The elemental composition of SOM includes carbon, oxygen, hydrogen, nitrogen, phosphorus, and sulfur. Nitrogen, phosphorus, and sulfur are plant nutrients that are slowly released during decomposition and are then available to plants, as well as other soil organisms. Other elemental nutrients may also be held in complex SOM.

1.2 Importance of SoilsSoil is a vital part of the natural environment. It is just as important as plants, animals, rocks, landforms and rivers. It influences the distribution of plant species and provides a habitat for a wide range of organisms.

Soils are important because they Filter water, Regulate the flow of water and chemical substances between the atmosphere and the earth, Provide habitat for millions of species of organisms, provide water, nutrients and support for plants, acts as both a source and store for gases (like oxygen and carbon dioxide) in the atmosphere e.g. it impounds carbon from the atmosphere, serve as a “compost bin” for the earth, moderate and regulate distributions of solar energy and provide a historical record of the influences of climate and living organisms upon the parent material from/in which the soil is formed.

pg. 4

2. SOIL FORMATION AND DEVELOPMENT.Soil formation (pedogenesis) is a complex process that does not follow a linear path of development. While some soils do form in place, most soils form in mineral or organic substrates (parent materials) that have been displaced from their source of origin.

2.1 Soil Formation FactorsThere are basically five factors that control the formation of soil, namely:

i) Parent Materialii) Climateiii) Topographyiv) Biotic Activity (contribution of Living organisms and man)v) Time

It is helpful to consider four basic processes as a framework for understanding how the five soil forming factors affect soil development. These four processes are:

a) Additions [surface or subsurface]b) Removals [ditto]c) Transformations [chemical weathering, turnover]d) Translocations [secondary clay minerals, base cations]

By considering processes, we can better understand how the biotic activity [BA] factor or any other factor influences soil properties.

2.1.1 Soil Parent MaterialParent materials for soils are the mineral or organic substrates in which a soil forms. Parent materials are classified primarily by their mode of deposition; however, Organic parent materials are distinguished by their low mineral content.

Inorganic Parent material is the initial mineral substance that forms a soil. It may reside at the site of its origin or be transported from somewhere else to its current location. A soil formed from parent material found at the site of its origin is called a residual or sedentary soil. Bedrock weathering in place produces a stony, massive material called saprolite. As physical and some chemical weathering occur, the saprolite becomes denser than the underlying bedrock.

The texture and original rock structure remain, but the material is soft enough to dig with a hand shovel. As chemical weathering converts primary minerals to secondary minerals, particles are redistributed vertically. As material is both added and removed, a soil develops. A residual soil will retain many of its characteristics from underlying bedrock.

Soil texture, mineralogy, pH, and other characteristics may be a direct result of the saprolite below.

Material can be eroded from one place and transported to another where it becomes parent material for a soil at the new site. Often weathering occurs before the material is transported to the new site. In this case, the soil may have few features in common with the underlying rock. Transported material can bury an existing soil at the new site. Once a depositional episode is completed, time zero for the new soil formation begins. Several forces can supply energy for the transportation of parent material e.g. wind, water, and gravity.

2.1.2 ClimateClimate is thought to have the greatest effect on soil formation. It not only directly affects material translocation (leaching or erosion, for example) and transformation (weathering), but also indirectly influences the type and amount of vegetation supported by a soil. Precipitation is the main force in moving clay and organic matter from the surface to a depth within the profile. When a soil is at field capacity, the

pg. 5

addition of more water will result in drainage either downward or laterally. Drainage water carries with it dissolved and suspended clay particles that collect at a new location within the soil profile. As a result, soils often show an increase in clay with depth as wind erosion selectively removes clay (and organic matter) from surface horizons.

Temperature and moisture affect physical and chemical weathering. Diurnal and seasonal changes in temperature cause particles to expand and contract unevenly, breaking them apart. Heat and moisture are active agents of chemical weathering, the conversion of one mineral into another.

Climate affects the type and amount of vegetation in a region. A warm, humid climate produces the most vegetative growth; however, microbial decomposition is also rapid. The net effect is that tropical and subtropical soils are generally low in organic content. In contrast, organic matter tends to be highest in a cool damp environment where decomposition is slow.

Temperature and the amount of water moving through a profile affects all of the following:

the amount and characteristics of organic matter; the depth at which clay accumulates; the type of minerals present; soil pH (humid climates tend to produce more acidic soil than do arid climates); soil color; iron, aluminum, and phosphorus distributions within a soil profile; and The depth to calcium carbonate and/or salt accumulation.

2.1.3 Topography (Relief)Topographic relief, or the slope and aspect of the land, has a strong influence on the distribution of soils on a landscape. Position on a slope influences the soil depth through differences in accumulation of erosional debris. Slope affects the amount of precipitation that infiltrates into soil versus that which runs off the surface. Aspect, or the direction a slope is facing, affects soil temperature. In northern hemisphere sites, south-facing slopes are warmer than those facing north. Differences in moisture and temperature regimes create microclimates that result in vegetational differences with aspect. Differences in weathering, erosion, leaching, and secondary mineral formation also can be associated with relief. Biotic Activity (contribution of Living organisms and man)

2.1.4 Biotic ActivityBiological activity and climate are active forces in soil formation. Soil pedogenesis involves a variety of animals, plants, and microorganisms. Ants, earthworms, and burrowing animals, for example, mix more soil than do humans through plowing and construction. Plant roots remove mineral nutrients from subsoil and redeposit them at the surface in leaf litter. Growing roots open channels through soil where rainwater can wash clay and organic matter down along these channels. Soil microbes decompose plant and animal debris, releasing organic acids. This biochemical activity is the catalyst for a great deal of the oxidation/reduction and other chemical reactions in soil.

The distribution of organic matter in a forest soil is different from that in a grassland. The surface soils of forests tend to have concentrated organic matter, which quickly decreases with depth. Grassland soils tend to accumulate organic matter to a greater depth than do forest soils.

2.1.5 TimeSoils develop over time. Soil formation is a dynamic process, where a steady state is slowly approached but only rarely reached. The rate at which a soil forms is related more to the intensity of other soil forming factors than to chronological age.

Soil development begins with a parent material that has a surface layer altered by vegetation and weathering. For example, a young Coastal Plain soil has relatively uniform material throughout, and is altered only by a dark-stained surface layer that has been formed by vegetation. A more mature soil, on the other

pg. 6

hand, shows evidence of the removal and transport of surface-layer clay to a subsurface layer called the B horizon. In an even older soil, chemical weathering and leaching have removed silicon, causing a change in the suite of clay minerals. A senile soil is excessively weathered and dominated by very resistant iron and aluminum oxide minerals. The rate that a young Coastal Plain soil becomes a senile soil depends not on its chronological age but on how rapidly minerals are transported and transformed within the profile.

Human activity frequently alters the process of pedogenesis. Once human activity ends, soil formation can continue as before if no radical change in the soil-forming factors occurred in the interim. Because fine material leaches selectively faster than coarse material, differences between human-altered and undisturbed soils in the ratio of fine to coarse clay may be apparent in a relatively short span of time (one hundred years in a humid environment).

2.2 Formation of Inorganic Soils

2.2.1 RocksA rock is defined by its mineral components. All mineral soils develop from parent rock. Soils developed from granite are obviously quite different from a piece of granite. Furthermore, all soils developed from granite are not the same. There should be some changes that take place in the composition of the rock during development of soil.

A rock is composed of one or more minerals. On the basis of origin, rocks may be divided into three kinds, namely igneous, sedimentary and metamorphic.

Igneous rocks arise from the cooling of molten magma. In East Africa, granite is an example of igneous rocks.

Sedimentary rocks are formed by consolidation of materials settled out of water. In east Africa, phyllite is an example of sedimentary rocks.

Metamorphic rocks are rocks altered in their form by heat and pressure usually during certain geological activities. In east Africa, schist and gneiss are examples of metamorphic rocks.

Granite is composed of quartz, feldspar and small quantities of dark-coloured ferromagnesium minerals. Since granite contains an abundance of silica, it is an acidic rock. On the other heart, basalt contains lower in silica and higher in bases, is a basic rock.

Minerals have definite chemical composition and are categorized largely on the basis of their crystalline structure. The group of minerals called feldspars have similar crystalline structures but differ in the bases which are present; thus a calcium feldspar is called anorthite, a potassium feldspar is called orthoclase, a sodium feldspar is called plagioclase, etc.

2.2.2 Physical weatheringThe disintegration of rocks occurs during heating and cooling. As a rock i.e. heated by exposure to the sun during the day, the surface of the rock becomes warmer than the interior of the rock. Due to the high temperature, the expansion of the surface is greater than that of the interior of the rock. When the sun goes down, the opposite happens. The rock looses heat by radiation to the atmosphere. Under these conditions, the surface of the rock becomes cooler than the interior and consequently, contracts to a greater degree. This differential expansion and contraction causes the rock to develop planes of weakness, and thin sheets of the rock tend to “flake off”. This flaking off is called exfoliation.

The erosive action of water, wind and gravity causes small pieces of rock to be carried along, striking other rocks and leading to their disintegration.

Physical weathering is relatively slow and only causes reduction in size of rocks. This reduction in size of rocks provides a larger surface which can undergo chemical weathering more easily.

pg. 7

2.2.3 Chemical weathering Hydration:

This happens when water unites with a mineral. When hematite (an iron oxide), commonly found in soils, is hydrated, it becomes limonite. This process of hydration is illustrated by the equation below:

Many other soil minerals can become hydrated causing them to expand, which results in splitting and breaking of the associated rock. Thus, chemical hydration contributes to the physical disintegration of rocks.

Oxidation:This is the chemical composition of oxygen from the atmosphere with the other elements. Iron and sulphur, in particular, are often in the reduced form and hence are subject to oxidation. Like all chemical reactions, oxidation is favoured by high temperatures. At low temperatures where the mean temperatures are higher, the extent and rate of oxidation is greater than at the cooler, higher elevations.

Hydrolysis:Water is a universal compound which often enters into chemical reactions. When the water molecule (HOH) splits into H and OH and each part combines with other elements, hydrolysis has occurred. The H in water tends to replace most of the positive ions in many other compounds, especially when the compound contains strongly basic elements. The hydrolysis of silicate or aluminosilicate minerals is very common. This may be illustrated by the following equations:

In both cases, the base was changed from its original form to a soluble form, or in terms of plants, an unavailable state to an available state.

Solution:

By this process, water is also able to alter the chemical properties of rocks. Water moving over or through rocks will dissolve and carry with it soluble portions of rock. The water will also carry the soluble products of decomposition such as the KOH and Ca(OH)2 (just mentioned) along.

The effectiveness of water in decomposing minerals is enhanced by the presence of salts, bases and acids. The mere presence of water aids in initiating and sustaining chemical rock and soil weathering.

Carbonation:After a certain time of rock weathering, primitive plants can begin to grow. They begin to contribute organic material, which stimulates growth of seeds from higher plants. The respiration of plant roots and decomposition of organic matter result in an enrichment of the medium with carbondioxide. This carbondioxide reacts with water to produce carbonic acid which is more effective in dissolving the minerals than pure water. The combination of carbondioxide with basic minerals such as KOH and Ca(OH)2 to form carbonates and bicarbonates is what we call carbonation.

2.2.4 Inter-relationship between Physical and Chemical WeatheringAll weathering processes e.g. disintegration, hydration, oxidation, hydrolysis, solution, carbonation may occur simultaneously and are somewhat interdependent. Chemical weathering takes place almost entirely on the surface of the rock particles. The finer the rock is subdivided, the grater the surface area will be, and the

pg. 8

faster weathering will occur. Physical disintegration of rocks has an important influence on the rate of chemical decomposition.

2.2.5 Rate of weatheringMinerals weather at different rates. Quartz is highly resistant to weathering. This explains why soils developed from rocks high in quartz have sandy textures. The quarts from the rock remains as sand particles. Potassium feldspars are also somewhat resistant to weathering but less than quartz. Some dark minerals such as hornblende and augite, are more readily weathered. Thus, rocks consisting of a mixture of minerals do not weather uniformly but rather, more readily weathered minerals decompose first.

Minerals might be classified on the basis of the kind of weathering which occurs, into three groups: Quarts, feldspars and the ferromagnesium minerals. When a granitic rock weathers, most of the easily weathered minerals break down initially thereby freeing the large amount of quartz grains which weather very slowly. At first, these quartz grains are sharply angular as they were in the original rock. But even quartz is weathered to a certain degree and as time passes, the sharp edges and corners of the quartz grains are broken. Eventually the quartz grains (or sand, as they would be called in the soil) appear rounded and abraded as though they had been subjected to the erosive action of being carried in a stream.

During quartz weathering, only a small amount of soluble silica is released which may be carried from the soil in solution or may recombine with other elements to form new compounds.

The potassium feldspars are among the first minerals to be weathered. Potassium feldspar (orthoclase) is an important constituent of granite. When granite is exposed to weathering, the orthoclase is the first mineral to break down. The decomposition of orthoclase is representative of the decomposition of all the feldspars. Weathering of orthoclase requires water and carbondioxide, both of which are usually abundant in soil. When rain the weak carbonic acid formed by rainwater and carbondioxide comes in contact with the orthoclase, the hydrogen ion (H+) from the acid displaces the potassium ion (K+) from the orthoclase crystal. The potassium then combines with the carbonate to form potassium carbonate, soluble salt. When potassium is removed from the orthoclase, it causes disruption of thee orthoclase crystal. When the crystal reforms, it is no longer an orthoclase crystal but is now the beginning of the clay mineral crystal. In the reorganization of the crystal structure there is some excess silicon and oxygen which will appear in solution as soluble silica (SiO2). The entire reaction is shown in the following equation:

The secondary “kaolinite” mineral formed is the beginning of the development of this type of clay minerals. Originally, it is probably amorphous, but as it combines with other units, the characteristic crystalline structure develops. Potassium carbonate is soluble and may be carried away in the drainage water, or may be used by plants, or combine with other elements to form new compounds, or the K + may be held by the clay mineral. The electronegative charge on the clay attracts the positively charged potassium ion K+, or in some cases the K becomes part of the clay mineral structure. The silica produced appears as soluble silica or as very finely divided quartz in the colloidal size range.

Decomposition of other feldspars follows a similar pattern. Of course, calcium feldspar would produce calcium carbonate instead of potassium carbonate and the CaCO3 produced might form the mineral calcite and subsequently be found in limestone or marble rock. Some calcium of course would eventually be carried away by water to the sea. The same pattern would be followed by sodium feldspars, with Na2CO3 being formed.

To a certain extent, weathering of ferromagnesium silicates produces the same products of weathering as the feldspars, i.e. clay, soluble salts, and colloidal silica. However the presence of iron and magnesium in these minerals makes possible the formation of additional compounds. Ironmay be incorporated into some of

pg. 9

the clay minerals or may form iron carbonates. Usually, iron unites with oxygen to form haematite4, Fe2CO3, with a characteristic red colour. The presence of red colour in soils is usually due to haematite. Other iron

oxides formed from weathering of iron-containing minerals include goethite {FeO(OH)} which has a brownish

colour, and limonite {2Fe2O3.3H2O} which has a yellowish- brown colour.

Magnesium from minerals is also released in soluble form, sometimes as a carbonate. More often, it remains in newly formed minerals or becomes part of certain clay minerals.

In summary:Un-weathered rock is neither physically nor chemically suited for plant growth. Essential nutrients in rocks are in an insoluble form and hence unavailable to plants. As rock decomposes, essential nutrients may be released from the mineral in a soluble form. The soluble nutrients may be leached from the soil and lost or recombine with other elements and remain in the soil.

As a rock weathers, there is a gradual loss of total nutrients since rock constituents are made more soluble and hence, more subject to loss from the soil. However the content of the available nutrients increases because some bases and other elements become more soluble. If the process of weathering continues long enough, many nutrients are lost or converted to stable, insoluble forms and both the total and the available nutrients in the resulting soil become very low.

Under conditions of tropical weathering such as prevail in Africa, rocks and soils are very thoroughly weathered, i.e., the level of nutrients is generally very low and only the more highly weathered clay minerals such as kaolinite and the hydrous oxides prevail. (see figure below).

This degree of weathering imparts very important properties to the soil. The fertility status of east African soils is largely a result of very intense weathering the soils have experienced.

Figure 7 some possible routes of weathering of soil minerals.

4 The greek word haimatites means “bloodlike”

pg. 10

3. THE SOIL PROFILE

Figure 1 The Soil Profile

3.1 Master horizons and layers

The capital letters O. A, E, B. C and R represent the master horizons and layers of soils. The capital letters are the base symbols to which other characters are added to complete the designation. Most horizons and layers are given a single capital letter symbol, but some require two. Currently seven master horizons and layers are recognized. The master horizons and their subdivisions represent layers which show evidence of change and some layers which have not been changed.

O horizon- (O stands for organic) includes litter layer, and sometimes a humus layer. These materials are differentiated by the degree of decomposition of the organic matter. This horizon is often missing in cultivated soils and severely eroded soils.

A Horizon - The layer commonly referred to as topsoil; This layer most closely resembles the ideal soil and is found below the O horizon and above the E horizon. Seeds germinate and plant roots grow in this dark-colored layer. It is made up of humus (decomposed organic matter) mixed with mineral particles.

E Horizon - This eluviation (leaching) layer is stripped of much of its clay and sometimes staining agents as water drips through the soil (in the elluviation process), and is thus often lighter in color than the others. It is lower in organic matter than the A horizon. It is beneath the A Horizon and above the B Horizon and is made up mostly of sand and silt, having lost most of its minerals and clay as water drips through the soil (in the process of eluviation). [E stands for elluvial].

B Horizon - Also called the subsoil - is a zone of illuviation (has accumulated substances—clays, organic matter, iron and aluminum compounds and calcium carbonate that have been leached from overlying horizons when mineralized water drips down). This layer is beneath the E Horizon and above the C Horizon.

C Horizon - Also called regolith: the layer beneath the B Horizon and above the R Horizon. It consists of slightly broken-up bedrock. Plant roots do not penetrate into this layer; very little organic material is found in this layer.

R Horizon - The unweathered rock (bedrock) layer that is beneath all the other layers.

pg. 11

4. BASIC PHYSICAL PROPERTIESThe characteristics of soil that a grower can see or feel greatly affect how soil can be used for the growth of plants or for other activities. E.g. if the soil is loose, the roots can grow easily through it. The physical properties of soil in consideration here include soil texture, structure, colour, rots, density, etc and the most commonly studied ones are soil texture which can be determined by quantitative analyses and soil structure which is more described in qualitative terms. Texture really does not change over a period of only a couple of hundred years or so, but structure can be changed rapidly, especially through management practices.

4.1 Soil TextureThe relative combination of sand, silt, and clay in a soil defines its texture. The particle sizes (diameters) in each of these three soil separates ranges between specific limits as reflected in figure 2.

Figure 2 Different Schemes of Soil Separates Classification

It is relevant to know that soil texture is important in determining the nutrient-holding abilities of a soil. Along with soil structure (the arrangement of soil particles in aggregates), the texture of soil is also important to water-holding capacity, water movement, and the amount and movement of soil air in a given soil. All of this is important to the health and type of plants and other organisms that can exist in a particular soil.

Soil texture is highly correlated with a range of soil chemical and physical properties. Fine textured soils with high clay contents generally have higher nutrient and water holding capacities than do coarse textured soils. However, fine textured soils often do not have drainage characteristics that are ideal for plant growth, especially if the soil does not have good structure.

Once the percent by weight of sand, silt, and clay are known (or, rather, any two of them), the soil texture can be plotted on the triangular graph known as the soil textural triangle (figure 3). The region on the graph where the three particle size percentages meet is the soil's texture. Loam has been determined to be the texture best suited to the growth of most agricultural crops, having the optimum combination of heavy and light soil qualities.

4.1.1 Soil textural classes Soils rarely consist entirely of a single separate, but instead are a mixture. Textural classes are based on different combinations of sand, silt, and clay. The twelve basic textural classes in order of increasing proportions of the fine separates and with appropriate abbreviations are:

1. Sand (s) 2. Loamy sand (ls) 3. Sandy loam (sl) 4. Loam (l)

7. Sandy clay loam (scl)8. Clay loam (cl)9. Silty clay loam (sicl)10. Sandy clay (sc)

pg. 12

5. Silt loam (sil) 6. Silt (si)

11. Silty clay (sic)12. Clay (c)

The term loam refers to soils having moderate amount of sand, silt and clay. Thus, loamy soils have textural properties intermediate to the properties of the individual separates.

A sandy loam soil has a soil texture somewhat coarser than a loam and loamy sand soil has a texture somewhat finer than sand. Those textural classes with the term sand in the name are often modified to indicate the fineness of the sand. Very coarse sand, for example, will have different properties than very fine sand. The ranges of sand sizes used by the USDA are below:

Table 2.1. 1 Sizes Limits on Sand Separates used by the USDASand separate Particle diameter(mm)Very coarse sand 2.0-1.0coarse sand 1.0-0.5Medium sand 0.5-0.25Fine sand 0.25-0.10Very fine sand 0.10-0.05

Actually, soil texture describes the proportions of soil l particles less than 2 mm in diameter and larger than 2 m mm in diameter. The separates larger than 2 mm in diameter, or coarse fragments , are described using terms such as stony, cobbly, gravelly, slaty, cherty, and flaggy. Each term has a precise meaning reflecting the size, shape, and composition of the coarse fragments.

Figure 3.USDA Soil Textural Triangle. (source: Soil Taxonomy, 1975. US Government Printing Office. Wachington, DC)The soil textural triangle is used when lab data is available.

The term “Loam” is used to refer to a USDA soil textural class that has fairly equal amounts of sand silt in clay.

pg. 13

4.1.2 Determination of soil textureSoil texture is determined either in the laboratory or in the field.

a. Pipette method:

This uses stokes law to measure settling rates of the soil particles. It is very accurate, however, very time consuming.

b. Hydrometer method

This puts into application the specific gravity of the soil. It is cheap and fast, but it is not as accurate as the pipette method.

c. Field method (Feel method)

It is the most widely used method of determining soil texture. It is quick, cheap, and no lab facilities are needed. However, many human errors are committed during the process, and the method requires a lot of training. (clays cause lots of problems)

pg. 14

Fig. 2.1.3 Flow Chart for Determining Soil Texture

pg. 15

4.2 Soil StructureSoil structure is the shape that soil takes based on its physical and chemical properties. It is the overall arrangement and orientation of soil particles (separates) into definite and usually, recognizable shapes/ forms. It is a combination or arrangement of primary soil particles into secondary units or peds. It involves the geometric arrangement of the particles and inter-particle forces that act between them. Soil structure has both direct and indirect influence on soil productivity; it influences how soil behaves, including how water moves into (infiltration), across (runoff), and through (percolation) a given soil.

The degree of distinction may be related to parent material, activities of humus or activities of organisms.

The structures are categorized by their shapes, sizes and grades. These shapes are mainly spherical or rectilinear.

4.2.1 Spherical Peds.This generally describes the units in spheroidal structures.

Spheroidal structures:they take the shapes of spheres. When very porous, they are considered crumbly. When they are not very porous, they are considered granular.

4.2.2 Rectilinear Peds.These generally describe units with platy, blocky and prismatic structures.

Platy structuresThey are such that their vertical thickness is much smaller than the horizontal dimension. They look like plates.

Prismatic structuresThe vertical sizing here is much greater than the horizontal dimensions. There is a special type of prismatic structures with tops that are rounded; these are known as columnar structures.

Blocky structures.There is not much difference between their vertical and horizontal dimensions. If the peds have sharp edged, they are referred to as angular blocky. If they have rounded edges, they are referred to as sub-angular blocky.

Maintaining good soil structure is important for plant growth. Below is a figure with some common structural characterizations.

pg. 16

Figure 4 Soil Structure Types

4.2.3 The Structure GradeIt is described in relation to hoe the peds appear in a soil exposure and how stongly they hold together when the soil is handled. The 4 classes of grades can be:

Strong Medium Weak Structureless.

Structureless soils have no visible geometric structure. They are subdivided into single grain soils and massive structures.

Single grain soils:Here, individual particles are not associated with other particles. They are generally sandy.

pg. 17

Massive structures:Here, the particles are strongly bound to each other, but not in any geometrically recognized pattern. These are common in soil horizons that have undergone little or no formation, and in places where fine textured soils are compacted together when wet.

At times, the soils can have compound structures in which large peds have broken up into smaller peds of different shapes. E.g. a coarse prismatic structure may break up into a medium blocky structure.

4.2.4 Soil structure formationPhysical, chemical and biological processes interact to form soil structures. Soil particles like clay act separately in a dispersed system or they group together to form domains in a flocculated system. Domains in turn, group together to form aggregates, and the aggregates form peds.

Figure 5 Representation of Soil Formation

Flocculated systems form stronger structures than dispersed systems because of particle interaction. If the salt content of the soil (Na+, Ca2+, Mg2+, Cl-, SO4

2-, etc in soil solution) is high, as in the desert, the soil tends to be flocculated.

In most soils however, the salt content is not high enough to cause flocculation. In these soils, the kind of exchangeable cations affects flocculation. Sodium ions, Na+ and to a less extent, Magnesium ions Mg 2+, tend to disperse soils. Whereas calcium ions, Ca2+ and aluminium cationic Al3+ tend to flocculate soils. Also iron and aluminium oxides tend to flocculate soils.

Formation of granular structuresBiological processes are mainly responsible for the formation of granular structures; ants and termites secrete organic “glues” that hold the soil particles together. Earthworms ingest soil, mix it with organic compounds and excrete aggregated soil material. Plants also secrete organic materials which bind soil particles together. These organisms normally affect the top soil, thus, top soils usually have granular structures.

4.2.5 Soil structure stabilityThis describes the ability of soil to retain its arrangement of soil voids when exposed to different stresses and is one of the most important properties that affect a soil’s ability to store and transport water and agrichemicals. For example, compaction alters the soil structure and hydrology by increasing bulk density, breaking down aggregates, decreasing porosity, aeration and infiltration, and by increasing soil strength and run-off.

4.3 Soil ColourThe first impression we have when we look at bare earth is the soil colour. Earth materials with different colours may imply the presence of some particular components. Red, yellow, yellowish-red, greyish-brown, and pale red are all good descriptive colours of soil, but not very exact.

pg. 18

Soil colour is easily noted but does not greatly affect the soil. However, it is a good indicator of certain soil conditions, so crop growers can learn about the soil condition by its colour.

4.3.1 Colour as a Guide to Soil UseSoil colour can be a suitable guide to the suitability of the soil for various uses. White or light coloured soils usually have low fertility, either because they are leached or high in salts. Proper irrigation, fertilizer application or treatment can render these soils usable. A black to dark brown soil colour usually suggests staining with organic matter. Very dark top soils that are high in organic matter may be quite fertile. However, sub-soils may be darkened due to gleying- the high organic matter content may occur due to lack of oxygen to aid in decaying it. Red indicates the presence of oxidized iron and is normally found in well drained soils. In soils saturated for long periods of the year, oxides become reduced giving a grayish or bluish gray colour.

4.3.2 Soil colour descriptionSoil colour is described using the Munsell method. This is a method used in soil survey that relying on a system which gives a precise description of soil colour. The method is known as the Munsell system of colour notation. (see www.munsell.com for more information)

4.3.3 The Munsell system of colour notationThe Munsell system identifies each chip with three variables:

Hue: This is the dominant spectral colour. It is related to the wavelength of the light reflected by the soil particles. Common soil colours are white, gray, black, yellow, brown, red and their various mixtures.

Value: The lightness or darkness of the hue/ colour. It is the measure of the amount of light reflected. Moisture content affects the lightness or darkness of soil.

Chroma: This is the strength or purity of the colour. It indicates the difference between white, black or neutral colour.

[i.e. Hue= specific colour, Value = lightnes or darkness, chroma = colour intensity]

Figure 6. The system recognizes 5 dominant colours with 5 additional “half way” colours such as yellow-red. These are referred to as the Hues. The Value describes the brightness of the colour from black (0) to white (10) and 5 being neutral grey. Chroma is a measure of “purity” or strength of the colour on a scale from 0 (weak) to 12 (very strong) – although the later only applies to some colours.

pg. 19

NB: Soil colour is strongly affected by moisture. As such, it is important to specify colours as either wet, moist, or dry.

4.3.4 Munsell Soil Color ChartsSoil color is characterized by comparison to the Munsell Soil Color Charts, which contain several series of distinctively colored chips. Each page represents a different hue. The Munsell book normally has 15 pages, each with a number (10, 7.5, 5, or 2.5) followed by a letter or letters indicating red (R), yellow (Y), green (G), blue (B), or combinations of these. For example, the 10 Y/R page contains color chips yellow-red (Y/R) with more yellow than red (10).Value units range between 0 and 10. The numbers ascend vertically on the page from the lowest to highest numbers, indicating dark to light values. Thus, a 0 value is black with no light reflected, while

10 is white with maximum light reflected. Chroma units are arranged horizontally across the page from 0 to 10, increasing in numbers from left to right. Low numbers indicate an increase in grayness, while high numbers signify a pure color with little mixing with other hues. Hence, a designation of 10R 6/4 indicates a hue of 10R, a value of 6, and a chroma of 4.

Mottling Mottles are spots or blotches that are different from the matrix or background colour.

Figure 7. soil with mottles

Figure 8: Gleying

Mottling is a blotchy colour condition in the soil that is indicative of less than perfect drainage but not as imperfectly or poorly drained that the blue/grey colours indicate. Mottles tend to be reddish in colour and stand out against the “matrix” colour of the soil.

On careful observation, most soils contain more than one color. Therefore, the matrix or dominant background color and mottles or colors different from the background must be described. While the matrix is simply described by a Munsell number, the mottles must be described by their abundance, size, and contrast to the background/matrix.

Abundance: the relative amount of mottling. It is described by three classes. Mottles that occupy less than 2 percent of the exposed horizon s are classified as few; 2 to 20 percent as common; and more than 20 percent as many.

pg. 20

Size: a measure of the estimated average diameter of individual mottles along their greatest dimension. Mottles less than 5 mm in diameter are classified as fine; 5 to 15 mm as medium; and greater than 15 mm as coarse.

Contrast: an indication of the relative difference in color between the matrix and mottles. If the contrast in color is only recognizable after close examination, it is classified as faint. A distinct pattern is readily seen although not striking. It may vary one or two hues or several value or chroma units. Mottles are considered prominent when they are the outstanding feature of the horizon. The colors of the matrix and mottles are separated by several units of hue, value, and chroma.

Gleying This is the extreme outcome or result of reduction by anaerobic bacterium. These bactirum do not use oxygen but a process called electron transfer respiration – they use the electrons from Fe2+ (ferrous iron) for energy and create Fe3+ (ferric iron) in the process. In this case the soil develops a blue grey colour. Soils that develop this colour are usually saturated for prolonged periods of time during the year.

4.4 Soil Density, ρAs already known, density is the mass per unit volume

4.4.1 Particle density ρs

Soil particle density is another way of expressing how much soil would weigh if there were no pore spaces. It varies according to the type of minerals from the parent rock and the amount of organic matter in the soil. Particle densities are very similar for most mineral soils. Most soils average about 2.6 gcm -3, a value used as a standard in soil density calculations. High organic matter reduces the value because organic matter is much lighter than mineral matter.

Mathematically,

4.4.2 Bulk density, ρb (dry density) This is the measure of soils compactness, defined as soils oven dried mass M s, divided by its volume Vb

including the pore space.

The bulk density of soil in good condition ranges from 0.8 to 1.6 gcm -3. Roots tend to proliferate more in soil with low bulk density.

Because soil has pores, bulk density (which is the actual density) is less than the particle density. To determine this, the soil is oven dried (at 150oC) and weighed to calculate the bulk density

Mathematically,

The bulk densities of mineral soils depend especially on the pore spaces in the soil. Organic soils are usually lighter, and soils with high organic matter tend to have good structure and lower bulk density than a similar soil with low organic matter. Cultivation destroys the structure by reducing organic matter and increasing the bulk density.

4.4.3 Soil Porosity, pThis is the total pore space in the soil, which is capable of holding the water and air; thus, 50% porosity is half solid particles and half pores.

pg. 21

It can be checked by placing an oven dry soil core in a pan of water until all the empty pores are filled with water. The water volume filling the pores divided by the total core volume is the porosity.

Porosity can also be calculated from the bulk density and particle density. If there were no pores, ρ b/ρs would equal to 1; so the ratio ρb/ρs is simply the soil percentage that is solid matter. If this is subtracted from 100%, it gives the percentage pore space. For calculation, it is usually assumed that the particle density, ρ s is 2.65 gcm-3

4.4.4 PermeabilityThis is the ease with which air, water and roots move within the soil. In highly permeable soil, water infiltrates so rapidly and aeration keeps roots well supplied with oxygen; roots also grow with ease.

Permeable soils are considered “loose” while impermeable soils are considered to be “tight” or “packed”

Permeability partly depends on the number of soil pores, but, mainly on the size and continuity of the soil pores.

Permeability is not a property that can be measured directly, however, the movement of water can be measured, and this reflects permeability. The measure of the rate of water movement in the soil is referred to as conductivity. Coarse textured soils have higher conductivities than fine textured soils.

NB: Bulk density can be related to soil particle density by the total porosity, according to the equation:

ρb = (1-pt)ρs

4.4.5 Total (wet density), ρt

When calculating this, we include the water component of the soil, so the soil weight is measured without oven drying. Total density differs from dry density in that it is strongly dependent on the Moisture content of the soil. If the soil is dry, ρt ≈ ρb

Mathematically,

; Mg ≈0

4.4.6 Void ratio, eThis is the volume of soil voids as a ratio of the solid volume; i.e. it is got by dividing volume of pores, Vf by the volume of mineral part, Vs

4.4.7 Soil wetness, θThis can be expressed on volume basis or on mass basis

pg. 22

4.5 Stages of Soil FormationAll soil formation begins with the accumulation of parent material. The next step is the buildup of organic materials at the surface. Pioneer species (most often grasses and alga in this area) live and die, and organic matter begins to build up on the surface of the material and also beneath the surface in the rooting zone.

The A horizon starts to form once enough organic matter has been transformed by soil biota into humic materials. The humic materials coat the inorganic soil particles, coloring them brown and black. The formation of a recognizable A horizon takes decades or, in some cases, centuries.

The B horizon begins to form as dissolved and suspended materials are carried downward to greater depths with percolating rainwater. These materials include humic substances, suspended clays, salts, and metals, including iron and aluminum. It is likely that the largely insoluble iron and aluminum cations and oxides move in complex with dissolved organic material (chelation), and also in complex with suspended clay minerals.

The A horizon continues to increase in thickness, and the B horizon continues to develop. The A horizon will increase in thickness and SOM content, until it reaches a steady state in which the rate of fresh organic matter additions equals the losses by decay, illuviation, and erosion. This steady state is affected by certain environmental changes, including climatic change and vegetational succession (or cultivation). The B horizon will continue to receive illuviated material as it is formed in the A horizon, or sometimes as it is deposited on the surface (especially wind-blown clays).

The E horizon forms as the top of B horizon moves deeper into the soil . In some forested areas, such as the Southeast region of the United States, the movement of illuvial materials occurs at a faster rate than the illuvial materials are formed (largely clays and organic matter). This results in a “gap” between the A horizon and the B horizon. The E horizon is usually the same texture as the A horizon, and the soil particles are largely stripped of staining agents, such as organic matter and metal oxides. These materials have elluviated from the E into the B horizon.

Minerals continue to weather. Clays in B horizon weather to less active minerals (kaolinite).

"Bases" are leached from soil. Certain cations are referred to as acids or bases in soil science, even though they do not fit any chemical definition of the term. The acidic cations, including aluminum and iron cations, are so called because their presence in the soil tends to decrease pH. (The reactions responsible for this will not be explained here.) The presence of the basic cations in large amounts usually coincides with neutral to high pH soil systems. These bases are often plant macro-nutrients, like calcium, potassium, and magnesium. The loss of basic cations results in low fertility soils.

Silicate clay minerals completely break down into iron and aluminum oxides. Soil is extremely infertile. This occurs in tropical climates.

pg. 23

5. CHEMICAL PROPERTIES OF SOIL

The chemical properties of soil usually bring into consideration the soil pH, salinity (EC), cation exchange capacity (CEC), organic matter and C:N ratio (Carbon to Nitrogen ratio)

5.1 Soil pH5:

Soil pH is a measurement of the acidity or alkalinity of a soil and is measured in pH units. The pH scale goes from 0 to 14 with pH 7.0 as the neutral point. As the amount of hydrogen ions in the soil increases the soil pH decreases thus becoming more acidic. From pH 7 to 0 the soil is increasingly more acidic and from pH 7 to 14 the soil is increasingly more alkaline or basic. A pH range of 6.8 to 7.2 is termed near neutral.

Descriptive terms commonly associated with certain ranges in soil pH are:

Extremely acid: < than 4.5; lemon=2.5; vinegar=3.0; stomach acid=2.0; soda=2–4 Very strongly acid: 4.5–5.0; beer=4.5–5.0; tomatoes=4.5

Strongly acid: 5.1–5.5; carrots=5.0; asparagus=5.5; boric acid=5.2; cabbage=5.3

Moderately acid: 5.6–6.0; potatoes=5.6

Slightly acid: 6.1–6.5; salmon=6.2; cow's milk=6.5

Neutral: 6.6–7.3; saliva=6.6–7.3; blood=7.3; shrimp=7.0

Slightly alkaline: 7.4–7.8; eggs=7.6–7.8

Moderately alkaline: 7.9–8.4; sea water=8.2; sodium bicarbonate=8.4

Strongly alkaline: 8.5–9.0; borax=9.0

Very strongly alkaline: > than 9.1; milk of magnesia=10.5, ammonia=11.1; lime=12

5.1.1 Measuring Soil pH

Soil pH provides various clues about soil properties and is easily determined.

- The most accurate method of determining soil pH is by a pH meter. - A second method which is simple and easy but less accurate than using a pH meter, consists of using

certain indicators or dyes.

Many dyes change color with an increase or decrease of pH making it possible to estimate soil pH. In making a pH determination on soil, the sample is saturated with the dye for a few minutes and the color observed. This method is accurate enough for most purposes. Kits (pH) containing the necessary chemicals and color charts are available from garden stores.

There may be considerable variation in the soil pH from one spot in a field or lawn to another. To determine the average soil pH of a field or lawn it is necessary to collect soil from several locations and combine into one sample.

5 Soil pH is defined as the negative logarithm of the hydrogen ion concentration.pg. 24

5.1.2 How pH Affects Nutrients, Minerals and GrowthHow pH affects nutrients, minerals and growth

Soil pH is an important chemical property because it affects the availability of nutrients to plants and the activity of soil microorganisms. The influence of pH on nutrient activity is illustrated in the figure below:

Figure 8. Availability of nutrients based on soil pH

Many gardening books list the preferred pH for common plants (generally 6.0 to 7.2). For most plants, however, what is preferred and what is tolerated are not related. Most plants tolerate pH of 6.0 to 7.5.

- The effect of soil pH is great on the solubility of minerals or nutrients.

Fourteen of the seventeen essential plant nutrients are obtained from the soil. Before a nutrient can be used by plants it must be dissolved in the soil solution. Most minerals and nutrients are more soluble or available in acid soils than in neutral or slightly alkaline soils.

Phosphorus is never readily soluble in the soil but is most available in soil with a pH range centered around 6.5. Soils that are highly aidic (pH 4.0-5.0) can have high concentrations of soluble aluminum, iron and manganese which may be toxic to the growth of some plants. A pH range of approximately 6 to 7 promotes the most ready availability of plant nutrients.

But some plants, such as irish potatoes and conifer trees, tolerate strong acid soils and grow well. Also, some plants do well only in slightly acid to moderately alkaline soils. However, a slightly alkaline (pH 7.4-7.8) or higher pH soil can cause a problem with the availability of iron causing chlorosis of the leaves (deficiency in chlorophyll) which will put the tree under stress leading to tree decline and eventual mortality.

- The soil pH can also influence plant growth by its effect on activity of beneficial microorganisms.

Bacteria that decompose soil organic matter are hindered in strong acid soils. This prevents organic matter from breaking down, resulting in an accumulation of organic matter and the tie up of nutrients, particularly nitrogen, that are held in the organic matter.

Table: Soil pH and Plant Growth

pg. 25

5.1.3 Changes in Soil pHSoils tend to become acidic as a result of:

(1) Rainwater leaching away basic ions (calcium, magnesium, potassium and sodium);

(2) Carbon dioxide from decomposing organic matter and root respiration dissolving in soil water to form a weak organic acid;

(3) Formation of strong organic and inorganic acids, such as nitric and sulfuric acid, from decaying organic matter and oxidation of ammonium and sulfur fertilizers. Strongly acid soils are usually the result of the action of these strong organic and inorganic acids.

Raising pH on acid soilsOn acid soils, the pH can be raised by adding lime (calcium carbonate). The addition of lime not only replaces hydrogen ions and raises soil pH, thereby eliminating most major problems associated with acid soils but it also provides two nutrients, calcium and magnesium to the soil. The amount to add depends on the cation exchange capacity (nutrient-holding capacity) of the soil, which is based on the soil’s clay content. Soil higher in clay will have a higher cation exchange capacity and will require more materials to raise the pH.

Lime also makes phosphorus that is added to the soil more available for plant growth and increases the availability of nitrogen by hastening the decomposition of organic matter. Liming materials are relatively inexpensive, comparatively mild to handle and leave no objectionable residues in the soil.

Some common liming materials are:

(i) Calcic limestone which is ground limestone; (ii) Dolomitic limestone from ground limestone high in magnesium; and (iii) Miscellaneous sources such as wood ashes.

The amount of lime to apply to correct a soil acidity problem is affected by a number of factors, including soil pH, texture (amount of sand, silt and clay), structure, and amount of organic matter. In addition to soil variables the crops or plants to be grown influence the amount of lime needed.

Lowering pH on alkaline soilsAssignment Question 1:

Explain how to lower pH of alkaline soils for agricultural aplications. ( 5 marks)

pg. 26

5.2 Salinity (EC)

The salinity of soil refers to the amount of salts in the soil and it can be estimated by measuring the electrical conductivity (EC) of an extracted soil solution.

5.2.1 How soil salinity affects plant growthSalinity can affect plant growth in several ways, directly and indirectly:The direct soil salinity damages include:

(i) Decreased water uptake;

High salts concentration results in high osmotic potential of the soil solution, so the plant has to use more energy to absorb water. Under extreme salinity conditions, plants may be unable to absorb water and will wilt, even when the surrounding soil is saturated.  

(ii) Ion-specific toxicity

When a plant absorbs water containing ions of harmful salts (e.g. Sodium, Chloride, excess of Boron etc.), visual symptoms might appear, such as stunted plant growth, small leaves, marginal necrosis of  leaves or fruit distortions.

The Indirect soil salinity damages include:

(i) Interference with uptake of essential nutrients

An imbalance in the salts content may result in a harmful competition between elements. This condition is called "antagonism", i.e. an excess of one ion limits the uptake of another ion. For example, excess of chloride reduces the uptake of nitrate, excess of phosphor reduces the uptake of manganese, and excess of potassium limits the uptake of calcium. 

(ii) Sodium effect on soil structure

In saline soils, sodium replaces calcium and magnesium, which are adsorbed to the surface of clay particles in the soil. Thus, aggregation of soil particles is reduced, and the soil will tend to disperse. When wet, a sodic soil tends to seal, its permeability is dramatically reduced, and thus water infiltration capacity is reduced as well. When dry, a sodic soil becomes hard has the tendency to crack. This may result in damages to roots.

It should be noted that salinity by itself actually improves soil structure and eliminates to some degree the negative effect of sodium ions, but of course, salinity cannot be increased without affecting plants growth.

5.2.2 Factors affecting soil salinity

Several factors affect the amount and composition of salts in soils namely: 

Irrigation water quality - The total amount of dissolved salts in the irrigation water, and their composition, influence the soil salinity. Therefore, various parameters, such as source water EC and its minerals content should be tested. 

Fertilizers applied - The type and amount of fertilizers applied to soil, affect its salinity. Some fertilizers contain high levels of potentially harmful salts, such as potassium chloride or ammonium sulphate. Overuse and misuse of fertilizers leads to salinity buildup, and should be avoided.

pg. 27

Irrigation regimen and type of irrigation system - The higher the water quantity applied, the closer soil salinity is to irrigation water salts concentration. When the soil dries, the concentration of salts in the soil solution is increased.

Since salts move with the wetting front, the salts accumulate in specific profiles according to the irrigation regimen and the type of irrigation used. For instance, when irrigating using sprinklers, water and salts move deeper, according to the soil's infiltration capacity and the water quantity, until they stop at a certain depth. When using drip irrigation - there is also a lateral movement of water and salts. 

Field's characteristics and agricultural history - A poorly drained soil might reach salinity level that is harmful to the plants and to the whole crop. A soil that was not flushed after a previous growing cycle might contain high level of accumulated salts. 

5.2.3 Prevention and Management of Soil Salinity

(i) Select a crop that fits the conditions in your fieldSoil type - water infiltration capacity, how much air does the soil contain, how much water will be needed to wash the soil in order to avoid salinity build up. Does your soil have special drainage problems? For example, it is better avoid planting a salt sensitive crop in a soil which is not well drained. 

(ii) The microclimate conditions in the field - Parameters such as wind direction and solar radiation may affect water consumption of the crop.

(iii) The agricultural history of the field Did salts accumulate in the soil during a previous crop? 

(iv) Irrigation water quality Check the quality of the available source water. What kind of salts does it contain and what is the total level of salts in it? 

(v) Type of irrigation system and its distribution What type of irrigation system are you going to use? Is it flood irrigation, sprinklers, pivot or drip irrigation? Each type of irrigation system has its own water distribution pattern, depending also on the soil properties. Make sure the emitters are set in the appropriate spacing, to allow uniform irrigation depending on your soil type.

(vi) Know the leaching requirement for your crop

Irrigation water amounts must coincide the growing stage of your crop. Apply the minimum needed to flush salts from soil. This means that you always have to give a little more water than the crop consumption, to allow leaching of salts below the root zone. Heavier soils require larger water applications than lighter soils, in order to avoid salinity buildup.

The leaching requirement is expressed as:

 

 A general equation to calculate the leaching requirement is

pg. 28

Where ECiw is the EC of the irrigation water, and EC th is the threshold salinity measured in the saturated soil extract, above which yield begins do decline (both in ds/m).

The total amount of water to be applied, AW, is

Where AW is the amount of water to be applied and ET is the water consumption based on evapotranspiration.

(vii) Keep the right Intervals between irrigations

Irrigation regimen and intervals must be appropriate to the soil conditions and to growth stage of the crop. Frequent and shallow (superficial) applications result in salt accumulation in the root zone, while larger applications, in longer intervals, will flush the salts below the root zone.

(viii) Use appropriate fertilizers types

The fertilizers type and their quantities should coincide with to the requirements of the crop and with nutrients which are already in the soil. There are fertilizers which contain salts which are not taken up by plants in large amounts, such as chlorides. These salts tend to accumulate in the soil.

(ix) Have your soil tested periodically

Soil analysis gives you a better indication of the salt content in the soil, without which you'll be only guessing. Guessing often comes close enough, but in many cases growers realize there's a salinity problem only after yields are decreased or crop quality is reduced.

A practical approach in order to prevent salinity buildup early enough is sampling the soil 5 times over a growing period of 8 months (a test every 6 weeks or so). It is recommended to do at least one water-analysis as well. The tests will indicate any change in soil content, allowing you to adjust the fertilization and irrigation regimen as needed.

This is the cheapest, most practical way to follow up on salinity status, keeping your crop quality and yield at optimal level.

(x) If problems persist

When you identify a salinity problem during the growing season, it is recommended to flush the field, even if it means risking some crop damage, rather than allowing further deterioration of the crop due to salinity.

Flushing applications should be carefully planned according to the crop conditions and growth stage. In light soils, which drain easily, the impact of flushing on the crop is usually insignificant.

pg. 29

In heavy soils, water infiltration and drainage problems may be encountered, resulting in excess of water and lack of air to the roots. Flushing heavy soils is a prolonged process and its final result is difficult to anticipate in advance.

Therefore, extra care should be taken when growing on heavy soils, as to not reach salinity buildup at all, or at least identify the problem early enough, when salts levels are still relatively easy to flush.

If all else fails and flushing is the chosen course of action, in heavier soils, not more than the maximal water amount that can be absorbed by the soil should be applied, and the longest intervals possible should be maintained. In the meantime, fertilization should be based only on Nitrogen and only the minimum amount should be applied.

The water used for flushing should be the highest quality possible, because the purpose of the flushing process is to decrease the soil salinity to the levels of the irrigation water.

5.3 Cation exchange capacity (CEC)Cation-exchange capacity is defined as the degree to which a soil can adsorb and exchange cations. Cation is a positively charged ion (NH4

+, K+, Ca2+, Fe2+, etc...) while Anion is a negatively charged ion (NO3-, PO4

2-, SO42-,

etc...)

Soil particles and organic matter have negative charges on their surfaces.  Mineral cations can adsorb to the negative surface charges or the inorganic and organic soil particles.  Once adsorbed, these minerals are not easily lost when the soil is leached by water and they also provide a nutrient reserve available to plant roots. These minerals can then be replaced or exchanged by other cations – this is what we call cation exchange.

CEC is highly dependent upon soil texture and organic matter content. In general, the more clay and organic matter in the soil, the higher the CEC.  Clay content is important because these small particles have a high ration of surface area to volume.  Different types of clays also vary in CEC.  Smectites have the highest CEC (80-100 millequivalents 100 g-1), followed by illites (15-40 meq 100 g-1) and kaolinites (3-15 meq 100 g-1).

Examples of CEC values for different soil textures are as follows:

Soil texture  CEC (meq/100g soi)Sands (light-colored) 3-5Sands (dark-colored) 10-20Loams 10-15Silt loams 15-25Clay and clay loams 20-50Organic soils 50-100

In general, the CEC of most soils increases with an increase in soil pH.  

Two factors determine the relative proportions of the different cations adsorbed by clays. First, cations are not held equally tight by the soil colloids. When the cations are present in equivalent amounts, the order of strength of adsorption is Al3+ > Ca2+ > Mg2+ > K+ = NH4+ > Na+.

Second, the relative concentrations of the cations in soil solution helps determine the degree of adsorption.   Very acid soils will have high concentrations of H+ and Al3+. In neutral to moderately alkaline soils, Ca2+ and Mg2+ dominate. Poorly drained arid soils may adsorb Na in very high quantities. 

5.3.1 Base SaturationThe proportion of CEC satisfied by basic cations (Ca, Mg, K, and Na) is termed percentage base saturation (BS%). This property is inversely related to soil acidity.  As the BS% increases, the pH increases. High base saturation is preferred but not essential for tree fruit production.  The availability of nutrient cations such as Ca, Mg, and K to plants increases with increasing BS%.

pg. 30

Base saturation is usually close to 100% in arid region soils.  Base saturation below 100% indicates that part of the CEC is occupied by hydrogen and/or aluminum ions.  Base saturation above 100% indicates that soluble salts or lime may be present, or that there is a procedural problem with the analysis.

5.3.2 CEC and Availability of NutrientsExchangeable cations, as mentioned above, may become available to plants.  Plant roots also possess cation exchange capacity.  Hydrogen ions from the root hairs and microorganisms may replace nutrient cations from the exchange complex on soil colloids.  The nutrient cations are then released into the soil solution where they can be taken up by the adsorptive surfaces of roots and soil organisms. They may however, be lost from the system by drainage water.

Additionally, high levels of one nutrient may influence uptake of another (antagonistic relationship).   For example, K uptake by plants is limited by high levels of Ca in some soils.   High levels of K can in turn, limit Mg uptake even if Mg levels in soil are high.

5.3.3 Anion ExchangeIn contrast to CEC, AEC is the degree to which a soil can adsorb and exchange anions.  AEC increases as soil pH decreases.  The pH of most productive soils in the US and Canada is usually too high (exceptions are for volcanic soils) for full development of AEC and thus it generally plays a minor role in supplying plants with anions.

Because the AEC of most agricultural soils is small compared to their CEC, mineral anions such as nitrate (NO3

- and Cl-) are repelled by the negative charge on soil colloids.  These ions remain mobile in the soil solution and thus are susceptible to leaching.

5.4 Organic matterSoil organic matter consists of a whole series of products which range from undecayed plant and animal tissues through ephemeral products of decomposition to fairly stable amorphous brown to black material bearing no trace of the anatomical structure of the material from which it was derived and it is the latter material that is normally defined as ‘humus’.

Humus is a complex and rather resistant mixture of brown or dark brown amorphous and colloidal organic substances modified from the original tissues or synthesized by the various soil organisms.

5.4.1 Non- humified & humified organic matter:Soil organic matter consists of two major types of compounds, unhumified substances and the humified remains of plant and animal tissues.

The non-humified organic matter is composed of compounds released during decomposition in the original or slightly modified form.

Although numerous organic compounds are present in the plant tissue, only a few exist in soils in detectable amounts after their release in soils. They are

Primarily

i. Carbohydrateii. amino acids and proteinsiii. lipidsiv. nucleic acidsv. lignin andvi. organic acids

Humified organic matter or humic matter is a group of compounds that includes humic acids, fulvic acids, hymotomelanic acid and humans. This humified soil organic fraction is also known as ‘humus’ or currently as

pg. 31

‘humic compounds’. Today humic compounds are defined as amorphous, colloidal poly dispersed substances with yellow to brown –black color and high molecular weights.

5.4.2 Sources of organic matter

5.5 Plant tissue The original/primary source of the soil organic matter is plant tissue. Under natural conditions, the tops & roots of trees, shrubs, grasses and other native plants annually supply large quantities of organic residues. Even with harvested crops, one tenth to one third of the plat tops commonly fall to the soil and remain there or are incorporated into the soil.

But these organic materials are decomposed and digested by soil organisms; they become part of the underlying soil by infiltration or by actual physical incorporation. Accordingly, the residues of higher plants provide food for soil organism, which in turn create stable compounds that help maintain the soil organic levels.

5.6 AnimalsThe secondary sources of organic matter are animals.

As they attack the original plant tissues, they contribute waste products and leave their own bodies as their life cycles are consummated. Certain forms of animal life, especially the earthworms, termites, and ants also play an important role in the translocation of soil and plant residues.

5.7 Soil organismsAnother source of organic matter is soil organisms. It is of ecological significance.

5.7.1 Types of materials present in organic matter:There are three types of material in soil organic matter. These are:

(i) Fresh or undecomposed material: Materials in which the anatomical structure of the plant substances is still visible.

(ii) Partially decomposed materials: Soft green portions are decomposed in this stage. This decomposition depends upon the composition of the tissues.

(iii) Completely decomposed material: In this stage, the materials are completely decomposed, and are called ‘humus’.

The whole of the organic residue is not decomposed all at once or as a whole. Some of the constituents are decomposed very rapidly some less readily, and others very slowly. The speed with which the organic residues are decomposed depends on the nature and abundance of various constituents that make up the residues, and the environmental condition, viz, moisture supply, aeration, temperature and soil reaction, under which the microorganisms carry out their activities. As the original material is decomposed new ones are simultaneously synthesized in the forms of microbial bodies and tissues.

It is evident that soil organic matter consists of plant, animal and microbial residues of various stages of decomposition and contains, at any given time,

(i) original residues;(ii) Various products of their decomposition, both simple and complex; (iii) Newly synthesized microbial bodies, both dead and alive.

5.7.2 Composition of plant tissues:On a volume basis, the moisture content of plant residues is high; varying from 60 to 90% and the rest 10 to 40% is solid matter.

pg. 32

On a weight basis, 75% is H2O & 25% is solid matter, there are 11% C, 10% O, 2% H & the rest 2% is others (N, P, K, Mg, As, etc. i.e., ash).

On an elemental basis (number of atoms of the elements), hydrogen predominates. In representative plant residues, there are 8 hydrogen atoms for every 3.7 carbon atoms and 2.5 oxygen atoms (H: C: O = 8:3.7:2.5). These three elements dominate the bulk of organic tissue in the soil. Even though other elements are present only in small quantities, they play vital role in plant nutrition and in meeting microorganism body requirements.

The actual organic compounds in plant tissue are many and varied and they can be grouped into a small number of classes. Representative percentages as well as ranges are shown in the following figure:

Figure 6: Typical composition of representative green plant materialThe carbohydrates, which range in complexity from simple sugars and starches to cellulose, are usually the most prominent of the organic compounds found in plants.

Fats and oils, which are somewhat more complex than carbohydrates and less than lignins, are found primarily in seeds & leaf coatings.

Proteins contain- in addition to carbon, oxygen and hydrogen – about 16% nitrogen and smaller amounts of other essential elements such as sulfur, manganese, copper and iron. Proteins are primary sources of these essential elements. Simple proteins decompose easily while the complex crude proteins are more resistant to breakdown and release their nitrogen.

Lignins which are complex compounds with multiple type or phenol structures are components of plant cell walls. The content of lignin increases as plants mature and is especially high in woody tissues. Lignins and polyphenols are notoriously resistant to decomposition.

5.8 C:N ratio (Carbon to Nitrogen)The ratio of carbon to nitrogen is the carbon: nitrogen ratio or the C:N ratio. There is always more carbon than nitrogen in organic matter. It is usually written as C:N and is a single number, because it expresses how much more carbon than nitrogen there is. For example if the ratio is 20, this means that there are 20 grams of carbon for each gram of nitrogen in that kind of organic matter. If the ratio is 100, it means that there are 100 grams of carbon for each gram of nitrogen. So if the number is low it means that the amount of carbon is reasonably similar to the amount of nitrogen. If the ratio is a large number, it means that there is

pg. 33

considerably more carbon than nitrogen. The C:N ratio does not tell us what form the carbon and nitrogen are in, just how much is there.

The C:N ratio of the organic material added to the soil influences the rate of decomposition of organic matter and this results in the release (mineralisation) or immobilization of soil nitrogen.

If the added organic material contains more nitrogen in proportion to the carbon, then nitrogen is released into the soil from the decomposing organic material.

On the other hand, if the organic material has a less amount of nitrogen in relation to the carbon then the microorganisms will utilize the soil nitrogen for further decomposition and the soil nitrogen will be immobilized and will not be available.

5.8.1 Importance of the C:N ration in soil

The C:N ratio is important because of what happens when organic matter is incorporated into soils. First, the larger organisms like mites and soil animals break it into smaller pieces. Then the fungi and bacteria start to decompose it (they secrete enzymes to break up the chemical compounds it is made of). When the enzymes have disrupted the compounds, the bacteria and fungi can use some of the parts released in this process as nutrients. For example, if the enzyme is degrading a protein, the microbe would be able to use the carbon, nitrogen and sulphur (if there is some) for its own cell wall structure and cellular contents. Excess nutrients to the requirements of the microorganisms are available for other soil organisms or plants to use. The microorganisms can access nitrogen in soil more easily than plants can, so the plants sometimes miss out. This means that if there is not enough nitrogen for all the organisms, the plants will probably be nitrogen deficient and nitrogen addition will be needed to meet the requirements of the plant. This is why incorporating organic matter into soils can change the amount of nitrogen (and other nutrients) available to plants. Incorporating organic matter that has a high C:N ratio will probably cause some nitrogen deficiency in the crops/plants, at least in the short-term.

5.8.2 Calculation Procedures of Carbon Nitrogen Ratio Example: A farmer incorporated 2,560 kg of fresh organic matter into the soil. It contained 55% carbon with a carbon nitrogen ratio of 20:1. As this material decomposes, will there be a decrease (immobilization) or an increase (release) in soil nitrogen? And by how many kg? Let us calculate the soil nitrogen status.

Fresh organic matter added = 2,560 kg Carbon content of organic matter= 55% Carbon: nitrogen ratio = 20:1

Step I : The amount of carbon present in the organic matter added to the soil.

2,560 kg of fresh organic matter x 0.55 carbon content =1,408 kg of carbon in organic matter

Step II: The amount of nitrogen present in the organic matter added to the soil. Fresh organic matter contains 1,408 kg of carbon and its C:N ratio is 20:1 .

The nitrogen content of the organic matter is

= 70 kg

Step III: The amount of C and N required for the microbes to form new tissues

As fresh organic matter is decomposed, the microbes use 75% of the carbon for energy and, the remaining 25% of the carbon is used to form their new tissue.

To form their new tissue, microbes use nitrogen from soil or from the added organic matter. Microbes require 1 kg of N for every 8 kg of carbon as the C:N ratio of microbes is 8:1.

pg. 34

Carbon content in fresh organic matter is 1,408 kg. Of this one fourth is used for new tissue.

The amount of carbon used by microbes for forming new tissue is

= 352 kg

The amount of nitrogen required by microbes to form the new tissues:

Microbes require 1 kg of N for every 8 kg of carbon to form the new tissue.

Therefore the amount of Nitrogen required for microbes using 352 Kg of Carbon is

= 44 kg 8

Step IV

Finally we compare the two nitrogen values. Fresh organic matter contained 70 kg of N, and microbes used 44 kg of N to form new tissue.

The balance nitrogen available is = 70 - 44 = 26 kg

NB: A POSITIVE balance indicates that the N in organic matter was more than the nitrogen required for the microbes, and this excess amount of N will be released into the soil, could be available for plants.

Try this problem:

Farmer incorporated 2,630 kg of fresh organic matter into the soil. It contained 51 % carbon with a carbon nitrogen ratio of 28 : 1. As this material decomposes, will there be a decrease (immobilization) or an increase (release) in soil nitrogen? And by how many kgs? Calculate the soil nitrogen status?

pg. 35

6. THE COLLOIDAL NATURE OF SOIL

6.1 Soil ColloidsThese are the tiniest particles in soil; known to be made of very fine clay and highly decomposed humus particles. They are too small to be seen with a light microscope, but their clear images can be made with electron microscopes. The upper limit of their diameter is usually given as 0.0001mm though particles somewhat larger may react similarly but to a lesser extent. For comparison, it would take 254,000 of these particles, side by side, to extend 1 inch (2.54 cm). Because of the sizing, they have very large reactive surface area hence, they are very active chemically and physically; Colloids are the most chemically active fraction of the soil and are closely associated with many reactions involved in plant nutrition.

Since colloidal clay and humus particles have negatively charged and positively charged sites, nutrient ions that are essentially for plant growth are attracted to the colloidal surface of the opposite charge. (The positively charged ions are called cations and the negatively charged ions are called anions). They are weakly held as a reserve supply to plants to draw upon. Without the attraction between the ions and colloids, the leaching of ions deeper into the soil and beyond the reach of roots would be much greater in humid regions.

The nature of the colloidal systems is not only dependent on the colloids themselves, but also on the ions attracted to them. These attracted ions may be exchanged partly according to the specific ions that are dominant (bigger in quantity) in solution. This process is called ion exchange. In almost all soils except some in tropical regions, the negatively charged sites on colloidal surfaces are much more numerous than positive sites, so the usual process is cation

6.1.1 Classification of colloidsSoils contain colloids which can be roughly divided into two general groups- organic and inorganic.

Organic colloids are formed as a result of activities of soil organisms. All organisms contain certain colloidal materials which are essential for the life of the organisms. The biological nature of organic colloids makes them more variable and to a degree more subject to control by man than the inorganic colloids.

Inorganic colloids are basically the clay minerals. They comprise much greater portion of most soils than the organic colloids and is more known of their structure and genesis.

6.2 ClaysMineral particles such as feldspar grains from granite are made up of mostly three elements: Si, O and Al. Therefore, they are called aluminosilicates. Small feldspar particles slowly change into clay minerals by weathering. These clay minerals are also aluminosilicates but they differ from feldspar in two ways: the clay minerals have some water molecules in their structure, and they have a platy or layered structure. They tend to have flat laminated, crystalline, plate-like shapes; i.e. they tend to fit together like the pages in a book or like pieces of mica which tend to break apart along parallel planes. This flat shape tends to give clays even a greater surface area than they might otherwise have. In addition to the normal surface area of a clay, some kinds of clays also tend to have what is called “internal surfaces”, i.e. areas on the interior of the crystals which may exhibit the same properties of adsorption and reactivity as the external surface.

Each clay particle, a micelle, usually has an electronegative charge which attracts positively charged ions to the clays. This property is very important for adequate plant nutrition. Clay minerals may be divided into silicate clays and hydrous oxide clays.

6.2.1 Mineralogy of clayThe basic crystalline structure of feldspars and some other minerals consist of a silicon atom surrounded by four oxygen atoms to form a tetrahedral shape similar to a 4 sided pyramid. These tetrahedral are bound

pg. 36

together to form sheets. The oxygen atoms of the adjacent tetrahedral are bound together through covalent bonds. These sheets of tetrahedral form an important part of the crystalline structure of clay minerals.

Another basic unit present in the clay minerals is the aluminium octahedron; this is made by six oxygen atoms surrounding one aluminium atom. The oxygen atoms are very large compared to the aluminium atom which occupies the space left in the centre of this arrangement. The planes created by this arrangement form an 8-sided figure, hence the name octahedron. The aluminium atom is larger than the silicon atom hence it requires larger space.

Atoms other than silicon and aluminium can exist in these spaces; size is an important factor in determining whether an atom can substitute in a crystalline structure.

Adjacent octahedral may share common oxygen atoms in a manner similar to the sharing of oxygen in the tetrahedral. Thus sheets of octahedral may be built up. In many cases, hydrogen is associated with the oxygen atoms producing hydroxyls. If all of the hydroxyls are shared, there are no excess charges and a

mineral called Gibbsite {Al(OH)3} results.

The silicon tetrahedral and the aluminium octahedral provide the basic building blocks for the clay minerals. The basic crystalline units are made up of combinations of the tetrahedral and octahedral. These are able to combine because the oxygens of the tetrahedral and the hydroxyls of the octahedral are nearly the same size. Thus, an apical oxygen in the tetrahedron may replace a hydroxyl in the octahedral layer. The substituted oxygen will then function in both layers; This may be repeated several times. With each substitution, some of the residual charge is dissipated resulting in a structure with a neutral charge.

Just as a plant leaf is made up of distinct layers of cells, the very small, flat clay crystals are made up of definite layers of ions. Most silicate clay particles are sandwich-like, with two silica layers (silicon plus oxygen) between which is an alumina layer (aluminum plus oxygen). They area called 2:1 clays because of this arrangement. Smectite and hydrous mica are clays of this type.

Clay of 1:1 type is like an open faced sandwich; there is a single silica layer adjacent to a single alumina layer ( Aluminium + hydroxyl). Kaolinite is a common 1:1 type clay. Plates of Hallosyte ( a variety of Kaolnite) tend to curl.

Figure 8 Clay particles are extremely small. In some cases the layers tend to curl

These 2:1 and 1:1 types of clays are called layer lattice silicate clays. The ions in each layer are arranged in a lattice-like geometric pattern.

Figure 9 ions in silicate clays tend to form a geometric pattern such as in this kaolinite

pg. 37

6.2.2 Groups of silicate claysSeveral groups of layer silicate clay minerals have been identified and within each group, there are many specific clay minerals. (only 3 are discussed to illustrate the nature & importance of clay)

Smectite groupMontmorillonite is a common member of the smectite group. It is a 2:1 type clay with a high capacity to hold plant nutrients and to swell and shrink on wetting and drying.

Each lattice consists of analuminium octahedra bound between two layers of silicon tetrahedral

Figure 10 Layer lattice crystals of montmorillonite clay.

Variations within this group are mainly due to the amount of substitution of Mg and ferrous ion for Al 3+ in the alumina layer. Soils that have high amounts of montmorillonite clay can be very troublesome particularly when wet. They are expanding lattice clays whereby their strong affinity for water causes the clay particles to spread apart and readily slip past one another. This results in what is called low bearing strength, which means that foundations of buildings and roads built on these clays are likely to fail (slip) and cause cracking in the super-structure, particularly on slopping ground. When montmorillonitic soils dry, cracks of nearly 2” (5cm) or more may open. Debris may fall into these cracks and cause the soil to buckle when it is wetted. Montmorillonitic soils become very sticky and difficult to till when wet and very hard when dry. As a result, farmers can work them only at just the right moisture content. Unimproved roads on montmorillonitic soils become impassable in rainy seasons.

The inner (alumina) layer of montmorillonite clay lattice is made up of Al, H and O ions. All the negative and positive charges balance and neutralize each other within this layer only if the three named ions are present. In montmorillonite clay, about a quarter of the Al3+

have been replaced by Mg2+ or Fe2+; ions with 2 positive charges have replaced ions with 3 positive charges. This produces a deficiency in positive charges which results in excess of negative charges at the surface of the crystal lattice. These are permanent negative charges that developed when the crystals were formed.

Smectite clays tend to be associated with the sub-humid and arid climatic regions. When found in humid regions, they are generally soils formed from shale or in the residue from basic rocks.

Hydrous mica groupHydrous mica has a rather slight structural difference from the primary mineral (mica) that is found in granite.

Figure 11 Layer lattice crystals of hydrous mica clays have lower capacity to hold plant nutrients and absorb water.

pg. 38

It is probably derived by the weathering of mica. It is associated with regions where weathering has not been severe and where the soil is neither very acidic nor basic. A member of this group is called illite. Hydrous mica is like montmorillonite in that it has a 2:1 lattice structure, but the lattice layers are held together by a mutual bond with potassium ions between them. This bonding minimizes the swelling and shrinking and results in good bearing strength for this clay and in reduced stickiness when wet. Illite has a lower capacity to hold plant nutrients than montmorillonite.

The presence of hydrous mica in a soil does not make the soil unstable in the way that montmorillonite does. A predominance of hydrous mica clay in a soil indicates lack of severe weathering. Such clays are likely to be found in the cool climatic zones where precipitation is high enough to remove soluble salts from the soil.

When the inter-layer K is removed completely by weathering, an expanding lattice 2:1 clay called vermiculite is formed. It does not shrink and swell as much as montmorillonite does. In vermiculite, the negative charge is derived from the isomorphous substitution of Al3+ for Si4+ in the outer layer. As a result, vermiculite has a higher negative charge than montmorillonite.

KaoliniteAn important clay mineral with a 1:1 lattice. Its basic structure consists of one layer of tetrahedral (silica) and one layer of octahedral (alumina).

Properties:Very low shrink-swell potentialCEC – 3-15 meq%Size: 0.1 – 5.0 micronsLow plasticity and cohesion.

Figure 12 Representation of Kaolinite group

Several of these basic units grouped together form a kaolinite particle; thus, a unit of kaolinite would be like a sheet of plywood and a sheet of soft-board glued together. The plywood would represent the tetrahedral layer and the soft-board would represent the octahedral layer, and the glue, the apical oxygens from the tetrahedral layer bonded to the octahedra. A stack of several of these units would correspond to a kaolinite particle.

It can be seen that kaolinite has the least silica of these silicate clays. This is the result of the intense weathering that is characteristic of warm regions of the world. One important property of kaolinite is the fixed spacing between the lattice layers. This is due to the attraction of hydrogen of the hydroxyl ions in an alumina layer for the oxygen in the adjacent silica layer. The bond between these lattice layers is of great importance because it renders kaolinite less sticky and gives the soil a greater bearing strength than with other types of silicate clays. Kaolinite has a very low capacity to hold plant nutrients and it absorbs less water than 2:1 clays.

6.2.3 Sources of negative charge on silicate clay mineralThe negative charge on silicate clays in soil comes from 2 sources. A typical silicate clay of the type 2:1 illustrates this principle.

First, the silica layer develops a negative charge from the oxygen ions along the edge of the crystal. Only one of the oxygen’s two negative charges is combined with the silicon ion, so at the plane where the crystal ends, there are oxygen ions with one negative charge unsatisfied. The figure below depicts this charge distribution in two dimensions, with an unsatisfied charge at each of the lattice.

pg. 39

Figure 13 development of negative charge on silicate clay

The oxygen ions are not shown in this schematic diagram, but their location is similar to the ionic arrangement in the silica layer shown in figure 9. This source of negative charge is called edge charge, although it is low, it is the main charge on the kaolinite clay. The edge charge in allophone is a bit higher. Edge charge is a pH-dependent charge because, under acidic conditions, the H+ ion can nullify the negative charge. Therefore, this is not permanent.

The second source of negative charge arises when one ion is substituted from another during the formation of the silicate clay crystal, without any change in its form. This is called Isomorphous (equal form) substitution, and it can occur in different ways. In some clays, the Al3+ ion substitutes for a Si4+ ion in the outer (silica layers), whereas in other clays a Mg2+ ion may substitute for Al3+ in the alumina layer. Either way, one negative charge results in the crystal and the charge is permanent.

6.2.4 Non-Crystalline Silicate ClaysWhen volcanic ash weathers in a relatively short time, some nearly amorphous silicate clays form. Two of the clay minerals are allophone, which is spherical, and imogolite, which is threadlike. Except for their small size, they share few of the properties of the layer silicate clays. This presence is germane to the classification of soils of volcanic origin.

6.2.5 Oxide ClaysApart from silicate clays, oxide clay and hydrated oxide clay minerals also exist in soils.

Figure 14 A particle of oxide clays.

Normally, these are oxides of Al and Fe, are amorphous, and are found most abundantly in soils formed from parent materials rich in Fe and AL in tropical and sub tropical regions where weathering has removed much of the silica from the clay fraction. Oxide clays have little or no crystallinity and very low capacity to hold plant nutrients. If Fe oxides are not very hydrated, they give soils a deep red colour. Ferrihydrite is one of the iron oxide clays that is very hydrated; it is important in the classification of some volcanic region soils.

Where Al oxide is concentrated enough in the sub-soil and below, it is called bauxite, an ore of Al. Bauxite is mined and loaded onto ships, and sent to areas where there is cheap electrical power for processing the ore to yield pure aluminum metal.

6.2.6 Cation ExchangeAlthough most soil colloids have a net negative charge, no electrical charge in the soil goes unbalanced for very long. Electrical neutrality is maintained.

A soil colloidal system has a double layer of charges. The inner layer is the negative charge of the colloidal layer discussed above. The outer layer is formed by cations in the soil solution, which are attracted to the colloidal surfaces in proportion to the negative charges available. This means that a divalent cation such as

pg. 40

Ca2+ or Mg2+ can neutralize two negative charges of the colloidal particle, whereas monovalent ions such as K+, Na+, H+ can neutralize one negative charge each.

Figure 15 double layer of charges in a colloidal system

In acid soils, Al3+ ions which may be combined with one or two hydroxyls(OH -), can be attracted to colloidal surfaces. There may be many other cations attracted to colloids in small amounts. Some of these are trace elements that are of great significance to growing plants.

Cations in the outer layer are sometimes called “swarm ions” because they resemble a swarm of bees around a hive, with the greatest concentration of the bees around the hive.

Figure 16 swarm of ions

In a soil colloidal system, these cations become hydrated so their effective radius includes the water molecules. The ions with two or more positive charges (such as calcium) with one positive charge and a large effective diameter to migrate farther from the surface of the colloid. All the attracted cations are in constant motion, but attraction holds them tightly enough so they are not easily lost to water that is moving through the soil. They are adsorbed ions because they are held to the surfaces of the colloids. This action is very important to plant life because it keeps many nutrients within the root zone of the crops.

Cations (e.g. Ca2+) in a mineral fragment are released by weathering into the soil solution where they are attracted to the particles of clay, around which they “swarm”. By exchange with hydrogen ions coming from around roots, the nutrient ions finally reach the roots. There is an area called oscillation zone in which ions are moving around roots and clay particles, this is the place of exchange, where one cation is replaced by another with an equal amount of charge.

Figure 17 A cation replaces another

For example, one divalent ion (such as Ca2+) may replace another divalent ion such as Mg2+or it may replace two mono-valent ions (such as two K+ ions). When a plant takes cations from the soil solution,

pg. 41

Figure 18 When a root takes in a cation into the plant.

When a plant takes cations from the soil solution, It releases hydrogen ions (H+) in exchange. For example when one Ca2+ is taken into the plant, two hydrogen ions are given off into the soil solution. Thus, electrical neutrality is maintained.

Cation exchange capacity

To quantify the negative charges on the soil colloids and therefore also the amount of cations attracted to those charges, it is essential to express the amount in standard units. The units are centimoles of charge per kg of soil material (cmolc/kg). the subscript-c indicates “charge”. The quantities determined are designated as Cation exchange capacity (CEC). Typically, this measurement is determined on soil samples, but it may be made on other earthly deposits such as lake-bottom sediments. There are many variations in the lab determination of CEC, but the basic principles behind the methods are similar:

(1) A known weight of soil is placed in a beaker and reacted with a solution containing only one type of cation, such as ammonium (NH4+).

(2) When it has been established that all the negative sites on the colloids have been satisfied with ammonium ions, the ammonium ions are replaced with other ions, and the ammonium ions replaced are measured.

(3) The cmolc /kg of NH4+ determined represents the CEC of the soil sample.

Frequently, the kinds of cations held on the colloidal system need to be determined. This can be done in a similar manner where in the exchangeable cations in the soil sample are replaced with another cation and those removed are annalysed individually. The kinds of cations found on the colloidal system of most soils are quite predictable. They are Ca2+. Mg2+, K+, Na+ and H+. in some soils, Al may also be very significant. When Nitrogen is added to the soil in form of NH3 or the NH4

+ ion, the adsorption of NH4+ becomes important. The

range in CEC for pure samples of the clays discussed here are tabulated below:

The range in CEC of some common clay materials

X CEC in cmolc/kgKaolinite 3 – 15 Illite 10 – 40 Montmorillonite 80 – 100 Vermiculite 100 – 150

For most agricultural soils, the CEC ranges between 3 to 20 cmolc/kg. The very sandy soils are at the low end of the scale and the very clayey soils or organic soils may have a CEC much higher than 20 cmol c/kg. The CEC of the soil gives a strong indication of the ability of the soil to retain and release nutrients, but it does not replace a soil test for plant nutrients.

pg. 42