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The Biology of SoilA community and ecosystem approach$

Richard Bardgett

Print publication date: 2005

Print ISBN-13: 9780198525035

Published to Oxford Scholarship Online: April 2010

DOI: 10.1093/acprof:oso/9780198525035.001.0001

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The soil environment

The soil environment

(p.1) 1 The soil environment
The Biology of Soil

Richard D. Bardgett

Oxford University Press

Abstract and Keywords

Soil forms a thin mantle over the Earth's surface and acts as the interface between the atmosphere and lithosphere, the outermost shell of the Earth. It is a multiphase system, consisting of mineral material, plant roots, water and gases, organic matter at various stages of decay, and a variety of live organisms. The first step towards understanding what controls the abundance and activities of these organisms, and also the factors that lead to spatial and temporal variability in soil biological communities, is to gain an understanding of the physical and chemical nature of the soil matrix in which they live. This chapter provides background on the factors responsible for regulating soil formation, and hence the variety of soils in the landscape. It also discusses the key properties of the soil environment that most influence soil biota, leading to variability in soil biological communities across different spatial and temporal scales.

Keywords:   soil formation, soil properties, soil-forming properties, soil texture, organic matter, soil pH

1.1 Introduction

Soil forms a thin mantle over the Earth’s surface and acts as the interface between the atmosphere and lithosphere, the outermost shell of the Earth. It is a multiphase system, consisting of mineral material, plant roots, water and gases, and organic matter at various stages of decay. The soil also provides a medium in which an astounding variety of organisms live. These organisms not only use the soil as a habitat and a source of energy, but also contribute to its formation, strongly influencing the soil’s physical and chemical properties and the nature of the vegetation that grows on it. Indeed, along with vegetation, the soil biota is one of five interactive soil-forming factors: parent material, climate, biota, relief, and time (Jenny 1941). The first step towards understanding what controls the abundance and activities of these organisms, and also the factors that lead to spatial and temporal variability in soil biological communities, is to gain an understanding of the physical and chemical nature of the soil matrix in which they live. This chapter provides background on the factors responsible for regulating soil formation, and hence the variety of soils in the landscape. It also discusses the key properties of the soil environment that most influence soil biota, leading to variability in soil biological communities across different spatial and temporal scales.

1.2 Soil formation

In order to understand the properties of soils that influence the biota that dwell therein, we must first consider some of the factors that lead to variations in soils and soil properties within the landscape. One of the most fascinating features of the terrestrial world is the tremendous variety in its landforms, reflecting a diversity of geological processes that have occurred (p.2) over millions of years; more recent as factors in the variation are biological processes and the influences of man. Similarly, within any landscape there is an incredible range of soils, resulting from almost infinite variation in soil-forming factors. These are highly interactive, in that they all play a part in the development of any particular soil. Combinations of these factors lead to the development of unique soil types, with a relatively predictable series of horizons (layers) that constitute the soil profile (Fig. 1.1). Of greatest interest to the soil ecologist are those horizons that are at, or close to, the soil surface; this is where most microbes and animals live and where most root growth and nutrient recycling occur. These horizons are referred to as the surface organic (O) horizon, which develops when decomposing organic matter accumulates on the soil surface, or the uppermost A horizon, which is composed largely of mineral material but also intermixed with organic matter derived from above. Soil ecologists are also concerned with the plant litter lying directly on the soil surface, deposited during the previous annual cycle of plant growth. This layer, referred to as the L layer, is often overlooked or even discarded in soil sampling regimes, but it is perhaps the most biologically active and functionally important zone of the soil profile.

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Fig. 1.1 Schematic representation of a soil profile showing major surface and subsurface horizons.


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Fig. 1.2 Podzol (Spodosol in US terminology) soil with deep O horizon (mor) and characteristic bleached Ea horizon above the red, depositional Bs or spodic horizon. (Image by Otto Ehrmann.)

Between this horizon and the A horizon are found layers of organic matter at intermediate stages of decomposition: the F layer, composed of partly decomposed litter from earlier years, and the H layer, made up of well decomposed litter, often mixed with mineral material from below.

While soil profiles vary greatly across landscapes, they can be classified into groups on the basis of their soil properties and soil-forming characteristics, each group having a unique set of ecosystem properties. For example, on free-draining, sandy parent material, in cold and wet climates, usually beneath coniferous forests, podzolic soils develop (Fig. 1.2). These soils, formed by podzolization (Box 1.1), have a deep, acidic surface O horizon, referred to as mor humus. They are subject to heavy leaching and are characterized by low rates of decomposition and plant nutrient availability, and hence low plant productivity. Typically the microbial biomass of these mor soils is dominated by fungi (rather than bacteria) and the fauna are characterized by high numbers of microarthropods (mites and Collembola), and an absence of earthworms. In contrast, on calcium-rich, clayey parent material, typically beneath grasslands and deciduous forests, brown earth soils are often found (Fig. 1.3). These soils have a mull humus composition that is often mildly acidic, owing to leaching of base cations (e.g. calcium) down the soil profile. The mull horizon is characterized by (p.4) (p.5) (p.6) intimate mixing of the surface organic and mineral-rich A horizon as a result of the high abundance and activity of soil biota, especially earthworms, leading to high rates of decomposition, nutrient availability, and plant growth. A total of 10 major soil groupings, termed soil orders, have been distinguished by the US Soil Taxonomy (Brady and Weil 1999) (Table 1.1), and 8 major soil groups are recognized by the Soil Survey of England and Wales (Avery 1980); each of these groupings has a unique set of ecosystem characteristics. Further details on these soils and their classification can be found in general soil science textbooks (White 1997; Brady and Weil 1999).

                   The soil environment

Fig. 1.3 A typical brown earth soil with mull horizon, under grassland. Note the lack of horizon differentiation caused by intense biological activity and the mixing of organic matter with mineral material from the A horizon (Image by Otto Ehrmann.).

Table 1.1 Soil taxonomy orders


Brief description


Recently formed azonal soils with no diagnostic horizons


Soils with swell-shrink clays and high base status


Slightly developed soils without contrasting horizons


Soils of arid regions


Soils with mull humus


Podzolic soils with iron and humus B horizons


Soils with a clay B horizon and >35% base saturation


Soils with a clay B horizon and <35% base saturation


Sesquioxide-rich, highly weathered soils


Organic hydromorphic soils (peats)

1.3 Soil-forming factors

As noted, within most landscapes there is a tremendous variety of soil types varying in physical and chemical make-up. The soil-forming factors are central to understanding the variability in soils at the landscape level and at the level of the individual soil profile. Being the central forces responsible for creating variety in soil conditions, and hence variations in the habitat of the soil biota, these factors require further consideration. The biota themselves, along with vegetation, constitute one of the main soil-forming factors; both can act as important determinants of soil formation and profile development. This section summarizes some of the important aspects of the main soil-forming factors. It is important to stress, however, that while soil-forming factors are considered individually, they operate interactively in nature, usually with a hierarchy of importance, with one or two of them being pre-eminent in soil development at a particular location.

(p.7) 1.3.1 Parent material

Geological processes acting over millions of years determine the variations and distribution of parent materials from which soils develop. Soils are formed from the weathering of either consolidated rock in situ or from unconsolidated deposits—derived from erosion of consolidated rock—that have been transported by water, ice, wind, or gravity. The mineralogical composition of these deposits varies tremendously. For example, the mineralogy of igneous rocks, formed by solidification of molten magma in, or on, the Earth’s crust, ranges from base-rich basalts (basic lava) with high amounts of calcium (Ca) and magnesium (Mg) to acidic rhyolites (acid lava) which contain high amounts of silica (Si) and low amounts of Ca and Mg. Rocks of intermediate base status, such as andesites, also commonly occur. Parent material also determines grain size, which determines soil texture (relative proportions of sand, silt, and clay), which in turn affects many soil properties, such as the ability of the soil to retain cations (its cation exchange capacity), the moisture retaining capacity, and soil profile drainage. Such variation in the mineralogy of rocks, therefore, strongly influences the type of soils that are formed and the character of the vegetation that they support (Fig. 1.4). Soils formed from weathering of basic lava, for example, tend to be rich in minerals such as Ca, Mg, and potassium (K) and fine textured (clayey), and have a high ability to retain cations of importance to plant nutrition (e.g. NH 4 + , Ca 2 + ). These soils are typically fertile brown earths with biologically active mull humus. In contrast, soils that are formed from acidic lava, such as granites and rhyolites, are low in Ca and Mg, coarse textured (sandy), and hence freely (p.8) drained, with low cation retention capacity. The soils that typically develop here are therefore strongly leached, nutrient-poor, acidic podzols with mor humus.

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Fig. 1.4 Schematic classification of igneous rocks and their resulting soils.

1.3.2 Climate

Historically, climate has been considered pre-eminent in soil formation, owing largely to the striking associations that exist, on continental scales, among regional climate, vegetation type, and associated soils. Indeed, these broad climatic associations led to the development in Russia of one of the first soil classification systems—the zonal concept of soils (Dokuchaev 1879). This system identified so-called zonal soils—those that are influenced over time more by regional climate than by any other soil-forming factor. While climate may play a crucial role in soil development on continental scales, for example, across Russia and Australia, it is arguably not as important in areas such as the subtropics and tropics where land surfaces are much older and more eroded, or in younger landscapes such as Britain where most soils are developed on recent, glacial deposits. Here, the factors of topography and parent material are of greater importance.

The effects of climate on soil development are largely due to temperature and precipitation, which vary considerably across climatic zones. These factors strongly govern the rates of chemical reactions and the growth and activities of biota in soil, which in turn affect the soil-forming processes of mineral weathering and decomposition of organic material. The effects of temperature on soil biological activity are well known; it is generally accepted that there is an approximate doubling of microbial activity and enzyme-catalysed reaction rates in soil for each 10 °C rise in temperature, up to around 30–35 °C. Above this temperature, however, most enzyme-catalysed reactions decline markedly, as proteins and membranes become denatured. Some microbes can live at extreme temperatures; for example, cold-tolerant fungi occur in polar soils and remain physiologically active down to −7 °C (Robinson and Wookey 1997). These cold-tolerant microbes are called pychrophiles, whereas microbes that live in extremely high temperatures are called thermophiles.

The effects of temperature and precipitation on soil formation are especially marked at high altitudes and latitudes. For example, soil organic matter content is often found to increase with increasing elevation, commonly reaching a peak in montane forests (Körner 1999). (At higher altitudes, above the treeline, the organic matter content of soils declines and reaches almost zero in unvegetated substrates in the upper alpine zone.) This increase in soil organic matter content is largely due to declines in temperature and high precipitation, which reduce microbial activity and rates of decomposition. Similarly, in high-latitude regions of Europe, vast peatlands have developed in areas where the combined effects of high (p.9) rainfall and low temperature, and minimal evapotranspiration, have led to anaerobic conditions (waterlogging) and the retardation of decomposition of organic matter. The consequence of this has been the accumulation of great masses of peat (blanket peats), especially in topographically uniform areas where drainage is reduced (Fig. 1.5). Dramatic changes in the physiology and productivity of dominant plants also occur along altitudinal and latitudinal gradients (Díaz et al. 1998), altering the nutritional quality (e.g. N content) of the leaf litter that is produced annually. As will be discussed in Chapter 4, such changes in organic matter quality resulting from shifts in plant community composition can have profound effects on the decomposability of organic matter, and hence the accumulation of organic matter on the soil surface.

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Fig. 1.5 Blanket peat at Moor House National Nature Reserve in northern England. Here, deep peats have developed at high altitudes where high rainfall and low evapotranspiration combine to cause excessive soil wetness that retards decomposition. These peatlands are of special significance because they represent a significant (ca. 30%) store of global terrestrial C. Indeed, the majority of the UK’s terrestrial C is stored in the peat soils of northern Britain (Image by Richard Bardgett.).

1.3.3 Topography

Variations in topography influence soil development, largely through effects on soil drainage and erosion. Soil drainage is primarily affected by the position of a soil on a slope; soils at or near the top of a slope tend to be freely (p.10) drained with a water table at some depth, whereas those at or near the bottom of the valley tend to be poorly drained with a water table close to the soil surface. These differences in drainage strongly influence soil development, leading to the development of a hydrological sequence (Fig. 1.6): welldrained soils on hilltops have deep, orange-brown subsurface horizons, indicative of oxidation processes (iron in ferric state); as drainage deteriorates towards the valley bottom, the soil profile becomes increasingly anaerobic and blue-grey in colour, indicative of a dominance of reduction processes (p.11) (iron in ferrous state) and the process of gleying (Box 1.1). In extremely wet valley bottom soils, deep O horizons develop on the soil surface as a result of retardation of decomposition processes under anaerobic conditions.

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Fig. 1.6 (a) Section of a slope and valley bottom showing a hydrological soil sequence, and (b) changes in soil profile morphology. (Redrawn with permission from Blackwell Science; White 1997)

Slope characteristics also greatly influence soil erosion processes, which in turn affect soil formation. In general, soils on ridges and steeper parts of slopes are shallower than those on lower slopes and valley bottoms, owing to the movement of soil particles down-slope by wash and soil creep. Because erosion preferentially moves finer particles down-slope, the soils of lower slopes and valley bottoms also differ in their mineralogical composition, being more fine textured. In humid regions, soils of lower slopes and valley bottoms also tend to be enriched in base cations and salts, owing to seepage of solutes from higher slopes and hilltops. Soil movement down-slope also leads to the formation of distinct morphological features on slopes, such as terraces. These features are especially common in alpine regions where steep slopes and freeze–thaw motion lead to instability of the surface soil and consequent down-slope creeping. This process is called solifluction. What happens here is that soils become saturated with water and freeze, and then melt; the expansion associated with freezing makes the surface soil very unstable when it thaws. This leads to downward movement of soil even on the gentlest slopes, especially if the subsurface soil is frozen.

1.3.4 Time

The age of the soil is a major factor underlying variations in soils and ecosystem properties. Soils become increasingly weathered over time, and, consequently, soil profiles generally become more differentiated, with more abundant and thicker horizons. This weathering process involves progressive leaching downwards of elements and minerals in percolating water. In particular, during the process of podzolization, progressive downward movement of iron (Fe) and aluminium (Al) leads to the development of a spodic (Box 1.1) subsurface horizon that is enriched with these minerals, and an upper Ea horizon, whence these elements have been removed, that is bleached in appearance. Over time, clay minerals also become leached down-profile, a process called lessivage (Box 1.1), leading to the development of subsurface argillic horizons. Other important changes in soils that occur with age are increases in soil organic matter and nitrogen (N) content (Crocker and Major 1955) and, over very long timescales (hundreds of thousands or millions of years), a progressive reduction in the availability of soil phosphorus (P) owing to its loss from the system and fixation in mineral forms that are not available to plants (Walker and Syers 1976).

The relationship between time and soil development is best illustrated by examining soil chronosequences, which are places where, for various reasons, a sequence of differently aged, but otherwise similar, geologic substrates exists. Glacier Bay, on the coast of southeast Alaska, is one of the (p.12) best known places for research on soil and ecosystem development because of the continuous retreat of the glaciers since 1794. Furthermore, records of the retreat, over some 100 km, have been maintained since this time, so the age of the glacial moraines is known. This has resulted in a site chronology over a period of 200 years, and from this, patterns of soil development can be tracked along with the succession in vegetation from the initial pioneer plant communities on recent moraine through to the climax spruce forest on the oldest moraines (Box 1.2). This chronosequence represents stages in the development of a podzol; as time progresses, the soil profile becomes progressively deeper and differentiated, the oldest soils being acidic in nature and having a thick organic surface horizon, a thin bleached horizon, and subsurface spodic horizon (Crocker and Major 1955). As organic matter steadily accumulates in the surface organic horizon, the amount of N in soil also increases; the organic carbon (C) and N content of underlying mineral soil also builds up (Fig. 1.7). Of particular significance for organic matter and N accumulation is the early stage of vigorous alder growth on terrain that has been ice-free for some 75 years. Here, N accumulates at a rate of about 4.9 g N m−2 yr−1, reaching values of some 250 g N m−2 within around 50 years of soil development (Crocker and Major 1955). Although there is no single explanation for the soil development sequence at Glacier Bay, it appears that establishment of plants and intensive leaching have played key roles (Matthews 1992).

Studies at Glacier Bay demonstrate progressive soil development over relatively short timescales towards climax, while other sites can be used to demonstrate soil change over hundreds of thousands, or even millions of years. The Hawaiian island archipelago, for example, presents a chronology of soil development over some 4.5 million years; Kilauea volcano on the Island of Hawaii is active, while Kauai, the oldest site, on the northwest end of the high islands is estimated to be 4.5 million years old (Fig. 1.8). A range of intermediate aged sites is also present, and all sites are derived from volcanic lava of similar mineralogy and have vegetation that is dominated by the same tree, Metrosideros polymorpha (Fig. 1.9). A key feature of this chronology is that it demonstrates a reduction in P availability in the oldest sites, which causes a dramatic decline in plant productivity (Crews et al. 1995). This fall in P availability follows the theoretical model of soil development that was proposed by Walker and Syers (1976): as soils age, P becomes surrounded, or occluded, by Fe and Al hydrous oxides, rendering the P largely unavailable to plants and also the soil biota (Fig. 1.10). This process of occlusion is especially likely to occur in very old soils because prolonged weathering of minerals leads to the formation of Fe and Al oxides that have a strong affinity for P. This P limitation to vegetation is further exacerbated in old soils because low soil fertility sets in motion a feedback whereby reductions in biological activity in soil reduce decomposition of plant litter, further intensifying nutrient limitation (p.13) (p.14) (p.15) (p.16) (Crews et al. 1995). Another excellent example of P limitation in very old sites is the large sand dune systems on the subtropical coast of Queensland, Australia. These dune systems provide a chronology of soil development, without interruption from glaciation, dating back some 400,000 years (Thompson 1981). Here, progressive weathering and leaching in the free-draining sand over this period has led to the development of so-called giant podzols (Thompson 1981) with soil profiles of some 20 m depth, which are characterized by extreme P limitation and stunted tree growth (Fig. 1.11). A detailed account of this soil development sequence can be found in Thompson (1992).

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Fig. 1.7 Total N content of soils recently uncovered by glacial retreat at Glacier Bay, Alaska. Plant succession is shown along the top. (Data from Crocker and Major 1955) (Reprinted with permission of Pearson Education; Kerbs 2001, using data from Crocker and Major 1955)

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Fig. 1.8 The Hawaiian Island archipelago presents a chronology of soil development of some 4.5 million years: Kilauea volcano on the Island of Hawaii is active, while Kauai, the oldest site, on the northwest end of the high islands is estimated to be 4.5 million years old. (Redrawn from Crews et al. 1995).

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Fig. 1.9 A Metrosideros polymorpha (‘ohi’a lehua) forest on the Hawaiian island of Molokai with low stature and biomass. This ecosystem developed in the absence of catastrophic disturbances during the past 1.4 million years. Depletion of P at this site has led to a decline in forest productivity. (Image by Richard Bardgett.)

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Fig. 1.10 Generalized effects of long-term weathering and soil development on the distribution and availability of P in soil (Adapted from Walker and Syers 1976).

1.3.5 Human influences

An increasing proportion of the Earth’s surface is under some form of management by man, and human impacts on ecosystems are significant and growing (Vitousek et al. 1997). The most obvious effects of humans on soil development result from the widespread destruction of natural vegetation for agricultural use, for example, in the tropics where some 8% of rainforest is thought to be cut down per decade (Watson 1999). Human activities also have dramatic effects on soil development by directly modifying the chemical and physical nature of the soil environment by fertilization, irrigation, and drainage, and by ploughing for cultivation. Humans can also have important indirect effects on soil development. For example, increasing concentrations of CO2 in the atmosphere have an effect on vegetation, (p.17) which in turn affects soil processes and soil biota. The introduction of invasive species into natural ecosystems can also have profound effects on soil properties and biota, for example, by changing the quality of litter inputs into soil and/or soil nutrient availability. Effects of human activities on soil biota and soil properties will be discussed in detail in Chapter 6.

                   The soil environment

Fig. 1.11 Stunted tree-growth resulting from P limitation in old-growth (400,000-year-old) eucalyptus forest on the subtropical coast of Queensland, Australia. (Image by Richard Bardgett.)

1.4 Soil properties

Variation in soil-forming factors determine the physical and chemical nature of soils, which in turn influences greatly the nature of the soil biota and hence ecosystem properties of decomposition and nutrient cycling. Variation in soil properties, especially the physical matrix of the soil, also greatly influences the movement of water and associated materials both (p.18) within and between ecosystems. This section examines some of the key soil properties that most strongly influence the soil biota and their activities, and the nature of ecosystems.

1.4.1 Soil texture and structure

The term soil texture refers to the relative proportions of various-sized particles—sand (0.05–2.0 mm), silt (0.002–0.05 mm), and clay (<0.002 mm)—within the soil matrix (Fig. 1.12). It is primarily dependent on the parent material from which the soil is formed and the rate at which it is weathered, as discussed above. Soil texture is of importance largely because it determines the ability of the soil to retain water and nutrients: clay minerals have a higher surface area to volume ratio than sand and silt, and hence soils with a high clay content are better able to hold water by adsorption and to retain cations on their charged surfaces. The ability of a soil to retain cations (e.g. Ca2+, Mg2+, NH 4 + ) is referred to as its cation exchange capacity, which reflects the capacity of clay minerals to hold cations on negatively charged surfaces. This retention of cations on clay minerals represents a major short-term store of nutrients for plant and microbial uptake.

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Fig. 1.12 Composition of the textural classes of soils based on percentages of sand, silt, and clay. For example, a soil with 60% sand, 10% silt, and 30% clay is a sandy clay loam.

(p.19) Soil structure reflects the binding of the above various-sized mineral particles into larger aggregates or a ped. The actual formation of stable aggregates requires the action of physical, chemical, and biological factors: (1) freeze–thaw and soil shrinkage and swelling help to mold the soil into aggregates; (2) the mechanical impact of rain and ploughing reorganizes soil materials; (3) the activities of burrowing animals, such as earthworms, lead to the mixing of mineral and organic materials and the formation of stable organo-mineral complexes; (4) the faeces of soil animals can act as nuclei for aggregate formation; (5) fine roots and microbes produce a range of polysaccharide glues which bind soil particles together, and fungal hyphae literally hold together mineral particles and organic matter. Together, these factors combine to produce stable aggregates in soil. As will be discussed in Chapter 3, large soil animals, such as earthworms and termites, can also substantially affect soil structure by creating macropores and channels as a consequence of their feeding and burrowing activities; this in turn increases infiltration rates, and hence drainage of water through soil.

Soil aggregation and structure is of concern to the soil ecologist, not only because the activity of the soil biota strongly affects it, but also because the structure of soil determines the physical nature of the living space. Aggregation determines the pore distribution of soil, which affects both the distribution of water in soil (specifically the degree to which pores are filled with water) and the extent to which biota are able to enter and occupy pore space, which is controlled by pore neck diameter and the size of the organism. For example, nematodes are approximately 30 μm in diameter, so their migration in soil is restricted by pores of diameter <30 μm and is optimal in soils with particle sizes in the region of 150–250 μm. Fungi and bacteria are much smaller in size, so they can occupy much smaller pores (1–6 μm) than can nematodes. Because the primary food source of many nematodes is microbes, these smaller pores can therefore provide a refuge for microbes from their predators, thereby affecting the nature and intensity of trophic interactions in soil.

In sum, good soil structure is recognized as a key attribute of fertile and biologically active soil, because it increases the flow of water and gases through soil, reducing the possibility of the development of anaerobic conditions, which would be detrimental to soil biota and their activities, and harmful to plant growth. Good soil structure promotes free movement of biota, thus increasing opportunities for trophic interactions; it allows roots to proliferate and enables aerobic microbial processes to dominate.

1.4.2 Soil organic matter

The organic matter content of soil varies tremendously in terms of both its chemical composition and its quantity. As already discussed, this depends (p.20) on a variety of interacting factors such as vegetation type, climate, parent material, soil drainage, and the activity of soil biota: a particular combination of these factors generally leads to the formation of either mull or mor humus forms. Soil organic matter is of importance because it promotes soil structural stability (see above), thereby preventing soil erosion. It is also extremely effective in retaining water within the soil matrix. Soil organic matter is of particular importance for biota because it is their primary source of nutrients and C. Through a range of activities, soil biota decompose organic matter, converting it into simple organic molecules that they can assimilate and use for energy and growth. Most components of soil organic matter occur as large molecules (e.g. cellulose, amino sugars, proteins, nucleic acid, lignin, etc.), which, over time, are depolymerized by extracellular microbial enzymes to yield simpler units which can then be assimilated by the microbes for energy and C. For example, the enzyme cellulase breaks down cellulose—the principal component of plant tissue—into glucose units that are readily assimilated by microbes and used for energy and C. Microbial enzymes also produce soluble organic nutrients (e.g. amino acids), which are either absorbed by microbes to meet their nutrient requirement (e.g. for C and N) or, in some circumstances, taken up directly by plants and mycorrhizal fungi (Chapter 3). Soluble organic nutrients that are absorbed by microbes are either retained to meet their C and N needs, or, when they require only C to support their energy needs, the excess N is secreted into the soil environment as inorganic N (ammonium), which is then available for plant uptake. This process is called mineralization; it is of key importance at the ecosystem scale because it determines the availability of inorganic nutrients for plant uptake, and hence plant productivity (Chapter 3). Soil fauna also contribute to the decomposition process and nutrient mineralization both directly, by mixing and fragmenting organic matter into smaller units that are more accessible to microbial attack, and indirectly, by feeding on microbes, affecting their growth and activity, and excreting nutrients that are in excess of the animals’ requirements into the soil environment (Chapter 3).

The rate at which organic inputs to soil are decomposed depends primarily on their quality, which is dependent on the type of compounds that are present within them. Rapidly decomposing materials, such as the litter of deciduous trees and animal faeces, generally contain high amounts of labile substances, such as amino acids and sugars, and low concentrations of recalcitrant compounds such as lignin. In contrast, the litter of coniferous trees decomposes slowly, being rich in large, complex structural compounds such as lignin and defence compounds such as polyphenols; this material is also unpalatable to soil fauna, further slowing down its decomposition. The importance of variation in the rate of decomposition at the ecosystem scale relates to the production of CO2 from heterotrophic microbial activity and its evolution into the atmosphere, and to the (p.21) conversion of organic nutrient forms to simple inorganic nutrients (e.g. ammonium and phosphate) that are available for plant uptake, a strong determinant of plant productivity. These factors will be discussed in detail in later chapters.

1.4.3 Soil water

It is water that renders the soil environment habitable. Together with dissolved nutrients it makes up the soil solution, an important medium for supplying nutrients and water to growing plants. It also provides a medium for many soil biota to live and move around in. Nematodes and protozoa, for example, live in water films and in free water, along with the bacteria that they feed upon. Larger fauna, such as microarthropods (mites and Collembola), live in open pore spaces but are very sensitive to desiccation, migrating out of soil when it becomes dry.

There are certain water-holding characteristics of soils that determine the amount of water that is available to plants and soil biota. Under normal conditions, soil pores will contain air as well as water, and most water is held in pore spaces as films or as water absorbed onto soil particles. Under these conditions the soil is said to be unsaturated. After a period of heavy rain or irrigation, however, pore spaces in soil become filled with water and the soil becomes saturated. After a period of drainage, when the amount of water held in the soil is in equilibrium with gravitational suction (5–10 kPa), field capacity is reached and no more water drains from the soil. The pores that are drained at field capacity are the macropores, which are those that are created by the burrowing activities of animals and by plant roots. As plant roots continue to absorb water from the soil, the amount of water held in films is reduced, and the remaining water adheres tightly to soil particles. At this stage, water moves through thin films to plant roots in response to a gradient of water potential; if water is not replenished and plants continue to take it up, a situation is reached when plants can no longer remove water that is strongly adhered to particle surfaces. This is the permanent wilting point, and the difference between this measure and field capacity is referred to as the water-holding capacity of the soil. In general, the water-holding capacity is greater in soils that contain large amounts of clay and/or organic matter, because these components have high surface areas that readily retain water. Also, clay soils have more small pores that readily retain water under gravitational suction than do sandy soils, which have larger pores that are more easily drained.

1.4.4 Soil pH

A large proportion of the Earth’s soils are acidic, especially in the tropics, where ecosystems persist at soil pH values of 4 or less (pH here is a measure (p.22) of the concentration of H+ ions in soil water). Many northern ecosystems also have very acidic soils: the pH values of the soils of Boreal forests and heathlands are often of 4 or less. In many parts of the world, soil acidity is further exacerbated by the use of inorganic fertilizers and acid rain (Kennedy 1992). Soils become acidic if base cations (e.g. Ca2+, Mg2+, K+) are leached from the soil profile, to be replaced by H+and Al3+ ions on cation exchange sites. Acidity in soils can come from various sources: (1) carbonic acid, which is formed by the dissolution of CO2 in water, dissociates to yield H+ ions; (2) microbial oxidation of ammonium ions (NH 4 + ) to nitrate (NO 3 - ) , the former being derived from mineralization and fertilizer inputs, also yields H+ ions; (3) atmospheric pollution (acid rain) and natural sources of acids, including volcanic eruptions and thunderstorms that yield sulphur dioxide and oxides of N, respectively, produce sulphuric and nitric acids that acidify soils (it has been proposed that the widespread occurrence of acidic soils in tropical and subtropical regions is, in part, a result of high thunderstorm activity in these regions. Long-term weathering and leaching of cations also contributes significantly to acidity in these old soils); and (4) decomposition of organic matter that has high concentrations of phenolic and carboxyl groups liberates H+ ions.

Soil pH is of concern to the soil ecologist because it controls nutrient availability and it directly impacts on soil biota. Acidic soils, for example, are characteristically high in soluble aluminium (Al3+), which can be toxic to plants and microbes. This is because aluminium occurs largely as insoluble forms, for example, as part of clay minerals whose structure becomes unstable at low pHs (4–5), releasing aluminium ions into soil solution (Kennedy 1992). The P availability in soil is typically low under acidic conditions, owing to the formation of iron and aluminium phosphates. These phosphates dissolve to release P into soil solution as pH rises, making it available for plant uptake; the availability of P is typically greatest between pH 6 and 7 (Chapter 3).

The effects of pH on soil organisms are well documented, and approximate tolerances of major groups of soil organisms are known. Most microbes grow within the pH range 4–9, although acidophiles are known to survive at pHs as low as 1, for example, sulphate-oxidizing bacteria (species of the genus Thiobacillus) which occur in hot springs and mine wastes. Growth in extremely alkaline conditions is restricted to a few fungi and bacteria, but such conditions tend to occur only in soda lakes and deserts, for example, those in Egypt which have pH values between 9 and 11. Soil animals are also very sensitive to acidic conditions in soil. For example, earthworms occur in very low numbers, with few species, in most acidic soils, and they become progressively more abundant as soil pH increases to neutrality (Edwards and Bohlen 1996). Acid soils tend to be dominated by enchytraeid worms (potworms), reaching densities as (p.23) high as 100 × 103 m−2 in acid peat soils (Cole et al. 2002a). Under such conditions, these enchytraeid worms replace earthworms as the functionally dominant soil animal.

1.5 Conclusions

Within most landscapes, there is a tremendous variety of soil types, reflecting spatial variations in the operation of a range of interacting soil-forming factors and pedogenic process, and the influences of man. The need for the soil ecologist to recognize this variation, at both spatial and temporal scales, cannot be stressed enough because it means that the roles of biota, relative to other factors, in controlling ecosystem processes (e.g. nutrient cycling, hydrological fluxes, and plant productivity) and soil formation itself will be context dependent. In other words, the roles of biota and the factors that regulate their community structure and activities will vary from soil to soil, depending on the dominant physical and chemical characteristics of that soil. In view of this, it is essential that studies on the soil biota are accompanied by detailed characterization of the physical and chemical nature of their habitat, and an appreciation of the soil-forming factors that led to the development of that soil in the first place.