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Chapter 4. Lithosphere as Microbial Habitat

Chapter 4. Lithosphere as Microbial Habitat

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FIGURE 4.1 (A) Pieces of granite showing phenocrysts, that is, visible crystals of mineral in a fine crystalline ground mass: an igneous rock. The inset fragment is 5 cm long. (B) Pahoehoe basalt from Kilauea,

Hawaii. Dark spots represent holes left by outgassing as the molten rock solidified. Note the absence of visible

crystals. (Courtesy of rock collection of the Department of Earth and Environmental Sciences, Rensselaer

Polytechnic Institute.)

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Lithosphere as Microbial Habitat



Minerals Classified Based on Mode of Formation

Primary minerals


Pyroxenes and amphiboles




Secondary minerals

Clay minerals




Hydrated iron and aluminum oxides


Source: Based on Lawton K. in Chemistry of the Soil, Van Nostrand Reinhold, New

York, 1955.




The mineral constituents of mineral soil are ultimately derived from rock that underwent weathering. Weathering, which leads to soil formation, is a process in which rock is eroded or broken down

into ever smaller particles and finally into constituent minerals. Some or all of these minerals may

become chemically altered. Some forms of rock weathering involve physical processes. For example, freezing and thawing of water in cracks and fissures of a rock may cause expansion of the cracks

and fissures and ultimate splitting of the rock because the ice formed from the water that originally

filled the cracks and fissures occupies a larger volume than the water from which it formed. Sand

carried by wind may cause sandblasting of rock surfaces. Alternate heating by the sun’s rays by day

and cooling at night may cause expansion and contraction of rock, leading to widening of cracks

and fissures. Waterborne abrasives or rock collisions may cause rock to break. Seismic activity may

cause rock to crumble. Evaporation of hard water in cracks and fissures of rock and resultant formation of crystals formed from the solutes in the hard water may cause rock to break because the

crystals occupy a larger volume than the original water solution from which they formed, thereby

widening the cracks and fissures through the pressure they exert. Mere alternate wetting and drying

may itself cause rock breakup.

Rock weathering processes may also be chemical when the weathering agents are of nonbiological origin. Examples are the solvent action of water; CO2 of volcanic origin; and mineral acids such

as H2SO3, HNO2, and HNO3 formed from gases of nonbiological origin, such as SO2, NO, and NO2,

respectively. Chemical weathering may also be caused by redox reagents of nonbiological origin,

such as H2S of volcanic origin or nitrate of atmospheric origin.

Finally, rock weathering may be the result of biological activity. Some of this activity may be

physical, as when roots of plants penetrate cracks and fissures in rock, forcing it apart. However,

much of it is biochemical, resulting from the activity of algae, fungi, lichens, and bacteria frequently

residing on rock surfaces and in the interior of porous rock. The microorganisms on the surface of

rocks may exist in biofilms, especially in a moist and wet environment. In biofilms containing a

mixed microbial population, the different organisms may arrange themselves in distinct zones where

conditions are most favorable for their existence (Costerton et al., 1994). Some microorganisms,

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the so-called boring organisms, may form cavities in limestone rock that they occupy by causing

dissolution of CaCO3 (Golubic et al., 1975). In other cases, opportunistic microorganisms invade

preformed cavities in rock (chasmolithic organisms) (Friedmann, 1982). Invertebrates, snails in

particular, may feed on boring organisms (Golubic and Schneider, 1979; Shachak et al., 1987) or

chasmolithic microorganisms by grinding away the superficial rock to expose the former to be consumed. The rock debris that the snails generate becomes part of a soil (Shachak et al., 1987; Jones

and Shachak, 1990).

Microbes dissolve rock minerals through the corrosive action of metabolic products, such as

NH3, HNO3, and CO2 (forming H2CO3 in water), and oxalic, citric, and gluconic acids they excrete.

Studies using scanning electron microscopy have shown that organic compounds formed by microorganisms such as lichens cause distinct weathering (Jones et al., 1981). Waksman and Starkey (1931)

cited the following reactions as examples of how microbes can affect weathering of minerals:

2KAlSi3O8 ϩ 2H 2O ϩ CO2 → H 4 Al 2Si 2O9 ϩ K 2CO3 ϩ 4SiO2



12MgFeSiO 4 ϩ 26H 2O ϩ 3O2 → 4H 4 Mg3Si2O9 ϩ 4SiO2 ϩ 6Fe 2O3 и 3H 2O





Reaction 4.1 is promoted by CO2 production in the metabolism of heterotrophic microorganisms,

and Reaction 4.2 by O2 production in oxygenic photosynthesis by cyanobacteria, algae, and lichens

inhabiting the surface of rocks. Further investigations have extended these observations. In recent

studies, reactions were examined in which organic acids that are excreted by microorganisms

promote weathering of primary minerals such as feldspars and secondary minerals such as clays

(Browne and Driscoll, 1992; Lucas et al., 1993; Hiebert and Bennett, 1992; Welch and Ullman,

1993; Brady and Carroll, 1994; Oelkers et al., 1994; Ullman et al., 1996; Bennett et al., 1996; Barker

and Banfield, 1996, 1998). Some current weathering models favor protonation as a means of displacing cationic components from the crystal lattice followed by cleaving of Si–O and Al–O bonds

(Berner et Holdren, 1977; Chou and Wollast, 1984). Others favor complexation, for instance, of Al

and Si in aluminosilicates, as a primary mechanism of dissolution (Wieland and Stumm, 1992;

Welch and Vandevivere, 1995).

Mineral soil may derive from aquatic sediment or alluvium left behind after the water that carried it from its place of origin to its final site of deposition has receded. Mineral soil can also

form in place as a result of progressive weathering of parent rock and subsequent differentiation of

weathering products. Soils originating by either mechanism undergo eluviation (removal of some

products by washing out) or alluviation (addition of new material by water transport). Any soil, once

formed, undergoes further gradual transformation due to the biological activity it supports (Buol

et al., 1980).


Mineral soil varies in composition, depending on the source of the parent material, the extent of

weathering, the amount of organic matter introduced into or generated in the soil, and the amount

of moisture it holds. Its texture is affected by the particle sizes of its inorganic constituents (stones,

>2 mm; sand grains, 0.05–2 mm; silt, 0.002–0.05 mm; clay particles, <0.002 mm), which determine

its porosity and thus its permeability to water and gases.

Many, but not all, mineral soils tend to be more or less obviously stratified. As many as three or

four major strata or horizons may be recognizable in agricultural and forest soil profiles. A soil profile is a vertical section through soil (Figure 4.2). The strata are labeled O, A, B, and C horizons. The

O horizon represents the litter zone, consisting of much undecomposed and partially decomposed

organic matter. Some soil profiles may lack an O horizon. The A and B horizons represent the true

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Litter zone










A horizon











CaCO 3


CaCO 3
















B horizon




































C horizon

FIGURE 4.2 Schematic representation of the major soil horizons of spodosol and mollisol. The litter zone is

also called the O horizon. The A and B horizons may be further subdivided on the basis of soil chemistry.

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soil. The C horizon represents the parent material from which the soil was formed. It may be the

bedrock or an earlier soil. The A and B horizons are often further subdivided, although these divisions are somewhat arbitrary. The A horizon is the biologically most active zone, containing most

of the root systems of plants growing on it and the microbes and other life forms that inhabit soil.

As is to be expected, the carbon content in this horizon is also greater. The biological activity in the

A horizon may cause solubilization of organic and inorganic matter, some or all of which, especially

the inorganic matter, is carried by soil water into the B horizon. At times, the A horizon is therefore

known as the leached layer,

r and the B horizon as the enriched layer.

r Both biological and abiological

factors play a role in soil profile formation.



Plants assist in soil evolution by contributing organic matter through excretions from their root

systems and as dead organic matter. The plant excretions may react directly with some soil mineral

constituents, or they may first be modified together with dead plant matter by microbes, resulting

in products that then react with soil mineral constituents. During their lifetime, plants remove some

minerals from soil and contribute to water movement through the soil by water absorption via their

roots and transpiration from their leaves. Their root system may also help prevent destruction of the

soil through wind and water erosion by anchoring it.

Burrowing invertebrates, from small mites to large earthworms, help to break up soil, keep it

porous, and redistribute organic matter. The habitat of some of these invertebrates is restricted to

specific regions in the soil profile.



Microbes contribute to soil evolution by mineralizing some or all of any added organic matter

during the decay process. Some of the metabolic products from this decay, such as organic and

inorganic acids, CO2, and NH3, interact slowly with soil minerals and cause their alteration or dissolution, which is an important step in soil profile formation (Berthelin, 1977; Welch and Ullman,

1993; Ullman et al., 1996; Barker and Banfield, 1996, 1998). For instance, the mineral chlorite has

been reported to be bacterially altered in this manner through loss of Fe and Mg and an increase

in Si. The mineral vermiculite has been reported to be bacterially altered through mobilization by

dissolution of Si, Al, Fe, and Mg; thereby forming montmorillonite (Berthelin and Boymond, 1978).

Certain microbes may interact directly (i.e., enzymatically) with certain inorganic soil minerals by

oxidizing or reducing them or their constituents (see Chapters 13, 14, and Chapters 16 through 19;

Ehrlich, 2001), resulting in their mobilization by dissolution, or in the formation of new minerals

(Berthelin, 1977). Microbes may also play an important role in humus formation.

Humus is an important constituent of soil, consisting of humic and fulvic acids, humins and

amino acids, lignin, amino sugars, and other compounds of biological origin (Paul and Clark, 1996,

pp. 148–152; Stevenson, 1994). Humic and fulvic acids are dispersible in solutions of NaOH or

sodium pyrophosphate whereas humin is not. Humic acids are precipitated at acid pH whereas

fulvic acids are not. The humus constituents humins and humic and fulvic acids represent components of soil organic matter that are only slowly decomposed. They are mostly formed by microbial

attack of plant organic matter introduced into the litter zone (O horizon) and in the A horizon.

Humus gives proper texture to soil and plays a significant role in regulating the availability of the

mineral elements that are important in plant nutrition and in detoxifying those that are harmful to

plants by complexing them. Humus also contributes to the water-holding capacity of soil. Some

microorganisms in soil can use humic substances as terminal electron acceptors in anaerobic oxidation of various other organic compounds and H2, and as electron shuttles in the anaerobic reduction

of Fe(III) oxides (Lovley et al., 1996, 1998; Newman and Kolter, 2000; Hernandez and Newman,

2001; Hernandez et al., 2004).

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Water from rain or melting snow may mobilize and transport some soluble soil components and

cause precipitation of others. This can contribute to horizon development as the water permeates

the soil. Precipitates, especially inorganic ones, that form in the soil water may lead to soil clumping

when formed in sufficient quantity. Water may also affect the distribution of soil gases by displacing

the rather insoluble ones, such as nitrogen and oxygen, and absorbing the more soluble and potentially corrosive ones, such as CO2, NH3, and H2S.



Only ∼50% of the volume of mineral soil is solid matter. The other 50% is pore space occupied by

water and gases such as CO2, N2, and O2. As might be expected, owing to the biological activity in

soil and the slow gas exchange with the external atmosphere, the CO2 concentration in the gas space

in soil usually exceeds that in air, whereas the O2 concentration is less than that in air. According

to Lebedev (see Kuznetsov et al., 1963), soil water may be distributed in distinct layers around soil

particles (Figures 4.3A and 4.3B). Surrounding a soil particle is hygroscopic water,

r a thin film of

Hygroscopic water





Pellicular water









FIGURE 4.3 Diagrammatic representation of soil water distribution according to Lebedev. (A) Water layers

around a soil particle when soil moisture is in excess of what the soil atmosphere can hold. (B) Movement

of pellicular water from soil particle (a), which is surrounded by it, to soil particle (b) lacking it; particles (a)

and (b) are in very close proximity. (Adapted from Kuznetsov SI, Ivanov MV, Lyalikova NN, Introduction to

Geological Microbiology, McGraw-Hill, New York, 1963.)

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3 ì 102 àm in thickness when surrounding a 25 mm diameter particle. This water never freezes

and never moves as a liquid. It is adsorbed by soil particles from water vapor in the atmosphere.

In a water-saturated atmosphere, pellicular waterr surrounds the hygroscopic water. Pellicular

water may move from one soil particle to another by intermolecular attraction, but not by gravity

(Figure 4.3B). It may contain dissolved salts, which may depress the freezing point to −1.5°C.

Gravitational waterr in Lebedev’s model surrounds pellicular water when moisture in excess of what

the soil atmosphere can hold is present. It moves by gravity and responds to hydrostatic pressure,

unlike hygroscopic water and pellicular water. So far, it is unclear as to which of these forms of

water is available to microorganisms. A reasonable guess is that gravitational water and probably

pellicular water can be used by them, but not hygroscopic water. The water requirement of microorganisms is usually studied in terms of moisture content, water activity, or water potential without

regard to the form of soil water (Dommergues and Mangenot, 1970).

Water activity of a soil sample is a measure of the degree of water saturation of the vapor phase

in the soil and is expressed in terms of relative humidity, but as a fractional number instead of percentage. Pure water has a water activity of 1. Except for extreme halophiles, bacteria have a higher

minimum water activity requirement (>0.85) than many fungi (>0.60) (Brock et al., 1984).

Water potentiall of soil is a measure of water availability in terms of the difference between the

free energy of the combined matrix and osmotic potentials of soil water and pure water at the same

temperature. Matric effects on water availability have to do with the effect of water adsorption to

solid surfaces such as soil particles, which lower water availability. Osmotic effects have to do with

the effect of dissolved solutes on water availability; their presence lowers it. As matric and osmotic

effects lower the free energy of water, water potential values are negative. The more negative the

water potential value, the lower the water availability. A zero potential is equivalent to pure water.

Osmotic water potential can be calculated from the effect of solute on the freezing point of water by

using the following formula of Lang (1967):

Water potential (J kg−1) = 1.322 × freezing point depression


where 100 J kg−1 is equal to 1 bar. Matric water requirements can be determined by the method of

Harris et al. (1970). In this method, NaCl or glycerol solutions of desired water potentials, solidified

with agar, are used to equilibrate with matrix material on which microbial growth is to occur. (For

further discussion on water potential, see Brock et al., 1984; Brown, 1976.)

The water potential requirement for two strains of the acidophilic iron oxidizer Acidihiobacillus

(formerly Thiobacillus)

s ferrooxidans has been determined by Brock (1975). Using NaCl as osmotic

agent, strain 57-5 exhibited a minimum water potential requirement at −18 to −32 bar, whereas

strain 59-1 exhibited it at −18 to −20 bar. Using glycerol as osmotic agent, strain 57-5 exhibited a

minimum water potential at −8.8 bar, whereas strain 59-1 exhibited it at −6 bar (Table 4.2). The

same study showed that significant amounts of CO2 were assimilated by A. ferrooxidans on coal

refuse material with water potentials between −8 and −29 bar, whereas none was assimilated when

the water potential of the refuse was less than −90 bar.


Organic or inorganic nutrients required by soil microbes are distributed between the soil solution and

the surface of mineral particles. Partitioning effects determine their relative concentrations in the two

phases. Their presence on the surface of soil particles may be the result of adsorption or ion exchange.

Nonionizable molecules tend to be adsorbed, whereas ionizable ones will bind as a result of charges

of opposite sign and may involve ion exchange. Microbial utilization of surface-bound molecules that

are metabolized intracellularly may require either their displacement from the mineral surface to be

taken into the cells or, in the case of some polymeric organic molecules, direct attack, for example,

by hydrolysis, of the portion of the molecule that is not bound to the surface. If displacement or direct

attack at the mineral surface cannot be effected, such nutrient will be unavailable.

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Effect of Osmotic Water Potential (Glycerol) on Growth of T. ferrooxidans


Total Water

Potential (bar)





















Strain 57-5





Strain 59-1




Note: All experiments were done in replicate tubes, which showed the same results. The incubation

period was 2 weeks. Iron concentration in the medium, 10 g of FeSO4 · 7H2O per liter; +, visible iron oxidation and microscopically visible growth; −, no iron oxidation nor microscopically visible growth.

Source: Reproduced from Brock TD, Appl. Microbiol., 29, 495–501, 1975. With permission.

Clay particles are especially important in ionic binding of organic or inorganic cationic solutes

(those having a positive charge). Such particles exhibit mostly negative charges except at their edges,

where positive charges may appear. Their capacity for ion exchange depends on their crystal structure. The partitioning of solutes between soil solution and mineral surfaces often results in greater

concentration of solutes on mineral surfaces than in the soil solution, and as a result the mineral surfaces may be the preferred habitat of soil microbes that require these solutes in more concentrated

form. However, ionically bound solutes on clay or other soil particles may be less available to soil

microbes because the microbes may not be able to dislodge them from the particle surface. In that

instance, soil solution may be the preferred habitat for microbes that have a requirement for such

solutes. Ionic binding to soil particles may be beneficial if a solute subject to such binding is toxic

and not readily dislodged (see Chapter 10).



Distinctive soil types may be identified by and correlated with climatic conditions and with the

vegetation they support (Bunting, 1967; Buol et al., 1980). Climatic conditions determine the kind

of vegetation that may develop. Thus, in the high northern latitudes, tundra soil, a type of inceptisol,

prevails, which in cold climate is often frozen and therefore supports only limited plant and microbial development. It has a poorly developed profile and may be slightly alkaline. Examples of tundra

soil are Arctic brown soil and bog soil. In the cool (i.e., temperate), humid zones at midlatitudes,

spodosols (Figures 4.2 and 4.4) prevail, which support extensive forests, particularly of the coniferous type. Spodosols tend to be acidic and have a strongly leached, grayish A horizon depleted in

colloids and compounds of iron and aluminum and a brown B horizon enriched in colloids and

compounds of iron and aluminum that are leached from the A horizon. In regions of moderate rainfall in temperate climates at midlatitudes, mollisols (Figures 4.2 and 4.5) prevail. These are soils

that support grasslands (i.e., they are prairie soils). They exhibit rich black topsoil and show lime

accumulation in the B horizon because they have neutral to alkaline pH. Oxisols are found at low

latitudes in tropical, humid climates. They are poorly zonated, highly weathered jungle soils with

a B horizon rich in sesquioxides or clays. Owing to the hot, humid climate conditions under which

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FIGURE 4.4 Soil profile of spodosol (podsol). (Courtesy of US Department of Agriculture (U.S.D.A.) Soil

Conservation Service, Washington, DC, USA.)

FIGURE 4.5 Soil profile of mollisol (chernozem). (Courtesy of US Department of Agriculture (U.S.D.A.)

Soil Conservation Service, Washington, DC, USA.)

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Lithosphere as Microbial Habitat


they exist, these soils are intensely active microbiologically and require constant replenishment of

organic matter by the vegetation growing on them and from animal excretions and remains to stay

fertile. The neutral to alkaline pH conditions of oxisols promote leaching of silicate and precipitation of iron and aluminum as sesquioxides. When oxisols are denuded of the arboreal vegetation, as

in slash-and-burn agriculture, they quickly lose their fertility as a result of intense microbial activity, which rapidly destroys soil organic matter. Because little organic matter is returned to the soil

in its agricultural exploitation, conditions favor laterization, a process in which iron and aluminum

oxides, silica, and carbonates are precipitated that cement the soil particles together and greatly

reduce the porosity and water-holding capacity of the soil and make it generally unfavorable for

plant growth.

Aridisols and entisols are desert soils that are present mostly in hot, arid climates at low latitudes.

Aridisols feature an ochreous surface soil and may show one or more subsurface horizons, such

as argillic horizon (a layer with silica and clay minerals dominating), cambic horizon (an altered,

light-colored layer, low in organic matter, with carbonates usually present), natric horizon (dominant presence of sodium in exchangeable cation fraction), salic horizon (enriched in water-soluble

salts), calcic horizon (secondarily enriched in CaCO3), gypsic horizon (secondarily enriched in

CaSO4 · 2H2O), and duripan horizon (primarily cemented by silica and secondarily by iron oxides

and carbonates) (Fuller, 1974; Buol et al., 1980). Entisols are poorly developed immature desert soils

without subsurface development. They may arise from recent alluvial deposits or from rock erosion

(Fuller, 1974; Buol et al., 1980).

Desert soils are not fertile due to the lack of sufficient moisture that prevents the development

of lush vegetation. However, insufficient nitrogen as major nutrient; and zinc, iron, and sometimes

copper, molybdenum, or manganese as minor nutrients may also limit plant growth. Desert soils

support a specially adapted macroflora and fauna that cope with the stressful conditions in such an

environment. They also harbor a characteristic microflora of bacteria, fungi, algae, and lichens.

Actinomycetes and lichens may sometimes be dominant. Cyanobacteria seem to be more important

in nitrogen fixation in desert soils than other bacteria. Desert soils can sometimes be converted to

productive agricultural soils by irrigation. Such watering often results in extensive solubilization of

salts from the subhorizons where they have accumulated during soil-forming episodes. As a consequence, the salt level in the available groundwater in the growth zone of the soil may increase to a

concentration that becomes inhibitory to plant growth. The drainage water from such irrigated soil

will also become increasingly salty and present a disposal and reuse problem.


Microorganisms found in mineral soil include prokaryotes (Bacteria and Archaea), fungi, protozoa,

and algae, and also viruses associated with these groups. Members of the prokaryotes may inhabit

the soil water and the surface of mineral soil particles. Some may be restricted to or be dominant in

soil water, whereas others may be restricted or dominant on the surface of soil particles. Prokaryotes

living on soil particles may form or inhabit biofilm, which may contribute to soil clumping. Mycelial

fungi may inhabit soil pores and spread on the surface of soil particles. Nonmycelial fungi, protozoa, and algae will dominate wherever their source of nutrients is most readily available. Viral

particles may exist in soil water or be adsorbed to soil particles.

A great variety of prokaryotes may be encountered in soil. A major portion of these have not

yet been cultured but are known to exist through analysis of DNA extracted from soil samples.

Morphological types of cultured Bacteria include gram-positive rods and cocci, gram-negative

rods and spirals, sheathed bacteria, stalked bacteria, mycelial bacteria (actinomycetes


s), budding

bacteria, and others. In terms of oxygen requirements, they may be aerobic, facultative, or anaerobic. Physiological types include cellulolytic, pectinolytic, saccharolytic, proteolytic, ammonifying, nitrifying, denitrifying, nitrogen-fixing, sulfate-reducing, iron-oxidizing and iron-reducing,

manganese-oxidizing and manganese-reducing, and other types. Morphologically the dominant

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forms of culturable Bacteria seem to be gram-positive cocci, probably representing the coccoid

phase of Arthrobacterr or possibly microaerophilic cocci related to Mycococcus (Casida, 1965). At

one time, non-spore-forming rods were held to be the dominant form. Spore-forming rods are not

very prevalent despite the fact that they are readily encountered when culturing the flora in soil in

the laboratory. Numerical dominance of a given type as determined by classical enumeration techniques employing selective culture methods does not necessarily speak of its biochemical importance in soil. Thus nitrifying and nitrogen-fixing bacteria, although less numerous than some other

physiological types, are of vital importance to the nitrogen cycle in soil. A given soil under a given

set of conditions will harbor an optimum number of individuals of each resident microbial group.

These numbers will change with modification in prevailing physical and chemical conditions. Total

viable counts of members of the domain Bacteria in soils generally range from 105 g−1 in poor soil to

108 g−1 in garden soil.

Certain members of the Bacteria in soil are primarily responsible for mineralization of organic

matter, nitrogen fixation, nitrification, denitrification (Campbell and Lees, 1967), and some other

geochemically important processes such as sulfate reduction, which cannot proceed in soil without

intervention of sulfate-reducing bacteria. Members of the Bacteria and certain Archaea play a significant role in mineral mobilization and immobilization. Some types of Bacteria, especially copiotrophs (microorganisms requiring a nutrient-rich environment), often reside in microcolonies or

biofilms on soil particles because the optimum nutritional and other requirements for their existence

are found there (see Section 4.2.6; Flemming et al., 2007). Conditions of nutrient supply, oxygen

supply, moisture availability, pH, and Eh may vary widely from one particle to another, owing in

part to the activity of different bacteria or other micro- or macroorganisms. Thus soil may contain

many different microenvironments. The colonization of soil particles by bacteria, especially those

forming exopolysaccharide (EPS), may cause some particles to adhere to one another (Martin and

Waksman, 1940, 1941), which means that bacteria can affect soil texture.

Fungi reside mainly in the O horizon and the upper A horizon of soil because they are, for the

most part, strict aerobes and find their richest food supply at these sites (Atlas and Bartha, 1997).

They are of great importance in the degradation of natural polymers such as cellulose and lignin,

which are the chief constituents of wood and which most kinds of bacteria are unable to attack.

Fungi share the degradation products from these polymers with bacteria, which then mineralize

them. Some fungi are predaceous and help control the protozoan (Alexander, 1977, p. 67) and nematode populations (Pramer, 1964) in soil. Their mycelial growth habit causes them to grow over soil

particles and penetrate the pore space of soil. They may also cause clumping of soil particles. The

soil fungi include members of all the major groups: Phycomycetes, Ascomycetes, Basidiomycetes,

and Deuteromycetes, and also slime molds. The last ones are usually classified separately from

fungi and protozoa, although possessing attributes of both. In numbers, the fungi represent a much

smaller fraction of the total microbial population in soil than bacteria. Total numbers of fungi,

expressed as propagules (spores, hyphae, hyphal fragments), may range from 104 to 106 g−1 of soil.

Protozoa are also found in soil. They inhabit mainly the upper portion of soil where their food

source (prey) is abundant. They are represented by flagellates (Mastigophora), amoebae (Sarcodina),

and ciliates (Ciliata). Like fungi, they are less numerous than the bacteria, typically ranging from

7 × 103 to 4 × 105 g−1 of soil (Alexander, 1977; Atlas and Bartha, 1997). The types and numbers

of protozoa in a given soil depend on soil type and soil condition. Although both saprozoic and

holozoic types occur, it is the latter that are of ecological importance in soil. Being predators, the

holozoic forms help to keep the bacteria, and, to a much lesser extent, other protozoa, fungi, and

algae in check (Paul and Clark, 1996).

In the study of soils, cyanobacteria and algae have been considered as a single group labeled

algae, although the cyanobacteria are prokaryotes and the algae are eukaryotes. Although both

groups are associated mostly with aquatic environments, they occur in significant numbers in the

O horizon and the uppermost portion of the A horizon in soils (Alexander, 1977; Atlas and Bartha,

1997) where sufficient light penetrates through translucent minerals and pore space. Overall, they

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