Jessica Gurevitch, Samuel M. Scheiner, and Gordon A. Fox - The Ecology of Plants - 2002 / The Ecology of Plants Jessica Gurevitch, Samuel M. Scheiner, and Gordon A. Fox; 2002 / Ch 04

gophers) animals. And, of course, soils are filled with the roots and rhizomes of plants. The living things in the soil continually alter its physical and chemical properties. Plants, for example, take up large volumes of water from the soil, changing the amount and distribution of water within the soil. The various soil organisms act and interact in a wide variety of ways. Some benefit one another; others consume one another or defend themselves to avoid being consumed. In the process of metabolism, soil organisms respire, altering the chemistry of their environment, and produce waste products and other substances, which become part of the soil. These activities affect plants in many ways, from making nutrients available to causing disease. We examine some of these effects in this chapter.

Soil Texture

The properties of soil, and its effects on plants, depend, first, on soil texture: the relative proportions of the different particles making up the soil. Soils are composed of sand, silt, and clay particles. Depending on which particle size dominates the character of a soil, the soil texture is categorized as sandy, silty, clayey, or loamy (Figure 4.1). The properties of loamy soils are a balance between sand, silt, and clay particles, and are generally considered the most desirable soils for agriculture. Sandy soils, with more than about 50% sand particles, have a coarse texture. These soils drain rapidly after a rain, and they hold water and minerals poorly. Water and air penetrate sandy soils easily. Because of these characteristics, they warm readily in spring and cool quickly in autumn.

Clay particles have very distinctive properties, and even a soil with as little as 35–40% clay will exhibit those properties (a little clay goes a long way). Clayey soils can hold a large volume of water, and they retain water and minerals exceptionally well. Similarly, clay-dominated soils retain pesticides, pollutants, and other substances. They are much less permeable to air and water than are sandy soils, which can result in puddling, greater runoff, poor drainage, and poor aeration (due to water filling up the pore spaces and excluding air). As a consequence, clayey soils are very slow to warm up in spring and cool more slowly in

Sand

Sand content (%)

Figure 4.1

(A) Examples of soil particle size of several different soil types. (B) The soil texture triangle, which shows how soil types are classified according to the percentage of sand, silt, and clay they contain (by dry weight). The values for clay are drawn parallel to the sand side of the triangle, those for silt are parallel to the clay side, and those for sand are parallel to the silt side. The points on the grid at which the three lines intersect defines the type of soil; for example, the grid lines for a soil with 20% clay, 40% sand, and 40% silt intersect in the region that shows a soil with this particular composition is loam.

(B)Kaolinite

Si tetrahedra

No water between lattice layers

Si tetrahedra

Montmorillonite

Si tetrahedra

Water and exchangeable cations between lattice layers

Si tetrahedra

14Å

Al octahedra (Isomorphous Mg)

Si tetrahedra

Illite

Si tetrahedra

Potassium ions between lattice layers

Si tetrahedra

10Å

Al octahedra (Isomorphous Fe, Mg)

Si tetrahedra

autumn; plant growth in these soils may be slow to get started early in the growing season. Silty soils tend to be intermediate in their characteristics and properties between sandy and clayey soils. Because soils often possess characteristics of more than one of these texture classes, they may be described by a combination of these terms, such as sandy loam, silty clay loam, or sandy clay.

To understand why different soil textures have the properties they do, we need to look more closely at the soil particles themselves. Sand and silt particles are generally irregular in shape, ranging from somewhat rectangular or blocky to chunky, spherical shapes (Figure 4.2A). In contrast, clay particles (Figure 4.2B) have a much more specific structure. These particles are made up of plates, whose size, shape, and arrangement depends on the minerals they contain and the conditions under which they were formed. A clay particle is made up of two or three flat crystalline plates, layered or laminated together. These particles can be hexagonal in shape with distinct edges, or they may form irregularly shaped flakes or even rods. Soils may also contain particles similar in size to clays, but without the distinctive crystalline structure of clay particles. For example, an amorphous material called allophane, with particles about the same tiny size as clay particles, is prevalent in soils developed from volcanic ash.

Most of the clay particles in soil exist in a colloidal state (sand and silt particles are much too large to form colloids). In a colloid, one or more materials in a finely divided state are suspended or dispersed throughout a second material. (Other examples of colloids are gelatin, fog, cytoplasm, blood, milk, and rubber.) Clay particles have an enormous amount of external surface area because they have a large surface-to-volume ratio and because there are so many of them in a given volume of soil. In some kinds of clays, there is additional internal surface area between the plates (see Figure 4.2B). Thus, the tremendous surface area that characterizes clays is due to the fineness of the particles and to their plate-like structure. The external surface area of a gram of fine clay is at least 1000 times that of a gram of coarse sand; the total surface area in the clay present in the

66 Chapter 4

top 10 cm of soil in less than a half hectare of clayey soil, if spread out, would cover the continental United States.

Because sand particles have a low surface-to-volume ratio, sandy soils have large, open pores between the mineral particles. Water drains easily from these pores because there is nothing to hold the water against the pull of gravity, and air penetrates them readily. Clay-domi- nated soils have a much larger number of pores than do sand-dominated soils. The total amount of pore space— the total proportion of the soil occupied by air and water—is greater in clayey soils (50–60%) than in sandy soils (35–50%), not only because of the smaller size of clay particles, but because of their arrangement. Clay particles (as well as particles of organic matter) tend to cluster together, forming porous aggregates, resulting in much more pore space than in sands, in which the particles lie close together. (Soils with native vegetation often have greater pore space than those that have been cultivated for crops because tilling damages some of the soil structure.)

The greater pore space is one factor contributing to the greater amount of water that can be held in claydominated soils, but there is another factor that is also very important. Unlike sand and silt particles, clay particles typically bear a strong negative electrochemical charge. Thus, they act as anions in the soil, attracting cations (positively charged ions, including the major plant nutrients), as well as water molecules, to their surfaces. The role of clay particles in adsorbing nutrient cations—attracting and holding them to their surface— is one of their most important effects for plants. Although there are many kinds of cations that are attracted to clay particles, certain ones are most prominent and most important for plant growth. In humid regions, hydrogen and calcium ions (H+ and Ca2+) are usually most abundant, followed by magnesium (Mg2+), potassium (K+), and sodium (Na+) ions. In soils of arid regions, hydrogen ions move to last place in this list, and sodium ions become more important. These positively charged ions attract numerous molecules of water to their surfaces, adding to the overall capacity of the clay particles to retain moisture in the soil.

The adsorbed cations on clay particles are partially available to be taken up by plants. There is a continuous, dynamic interaction between the ions adsorbed on clay and other colloidal particles and those in the soil solution— the water in the soil and its associated dissolved minerals. Ions that are displaced from their positions on the surface of a particle enter the soil solution, from which they can be taken up by plants, leached (lost from the surface soil as water drains), or adsorbed on another particle. These reactions differ greatly among soils of different origins, textures, and chemical compositions, and among regions differing in temperature and especially in rainfall.

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Soil pH

The pH of soil—the negative logarithm of the concentration of H+ ions in the soil solution—varies widely among different soils. The pH scale ranges from 1 to 14, where 7.0 is neutral (the pH of pure water); acid soils have lower pH numbers and higher H+ ion concentrations. Soils in the United States can range from less than 3.5 (for example, in the pine barrens of New Jersey) to as high as 10 (for example, in arid grasslands of the southwestern United States). For a soil to have a pH above 7 (neutral soil), it must be calcareous (containing CaCO3), sodic (containing Na2CO3), or dolomitic (containing CaCO3 • MgCO3). Most crops grow best in slightly acidic soils, but native vegetation can be adapted to anything from very acid to neutral to alkaline soils.

Soil pH has enormous effects on plant growth, and indeed, in determining what species can survive and grow in the soil. It acts on plants indirectly, however, through its strong effects on the availability of mineral nutrients and on the activity of some soil organisms (such as bacteria and fungi), changing the conditions for plant growth in a complex manner. Forest trees can grow over a range of soil pH, but are especially tolerant of acid soils. Conifers and some other tree species tend to increase the acidity of the soil in which they are growing, primarily through the properties of the litter they produce (shed needles, etc.). Grasslands tend to be found on relatively alkaline (high pH) soils, but this may be an indirect effect because low rainfall results in soils with high pH (as we will see below) and also favors grassdominated vegetation. There are other characteristic associations between soil pH and plants. Species in the Ericaceae, for example, such as heaths and blueberries, tend to grow only in very acid soils, while other taxa, such as Larrea tridentata (creosote bush, Zygophyllaceae), are typically found in alkaline soils.

Soil pH can strongly affect the availability of cations to plants. Cations bind loosely to clay particles, which are generally negatively charged. Under acid conditions, the excess H+ ions tend to bind more strongly to the clay particles than do the nutrient ions, displacing these nutrients into the soil solution. Thus mild acidity, which is characteristic of many soils, promotes nutrient availability. Under extreme acidity (whether natural or due to acid precipitation), however, the nutrient cations are so mobile that they are easily leached, and are carried off in groundwater—water found underground in aquifers, rock crevices, and so on.

What determines the pH of a soil? Two cations, hydrogen and aluminum, tend to increase soil acidity, while the other cations have the opposite effect, increasing soil alkalinity. Hydrogen ions are continually added to the soil by decaying organic matter, roots, and various soil organisms. They are actively exchanged between the surfaces of colloidal particles in the soil

(such as clay and organic matter) and the soil solution, contributing directly to the acidity of the soil. Aluminum ions (Al3+) indirectly cause hydrogen ions to be released from colloidal particles by reacting with water to form Al(OH)2+ and Al(OH)2+ plus H+. These hydrogen ions are then added to the H+ ions in the soil solution. High rainfall levels favor the predominance of aluminum and hydrogen ions because these ions are held very strongly by colloidal particles, while the other cations are more readily leached and thus lost from the soil.

Most of the other cations, called exchangeable bases, contribute to making the soil more alkaline. The cation bold exchange capacity (CEC) of a soil is a measure of the total ability of the soil colloids to adsorb cations (in units of centimoles of positive charge per kilogram of dry soil [cmolc/kg]), and the percentage base saturation is the proportion of the CEC that is occupied by exchangeable bases (Figure 4.3). In arid regions, the bases are not leached out of the soil by rainfall, so the percentage base saturation is very high (90–100%), H+ concentration is low, and the soils tend to be alkaline. In areas with higher rainfall, the bases are leached more easily while H+ and Al3+ ions are retained, so the percentage base saturation

is much lower (50–70%), and the soils tend to be acid. Plants and soil organisms are an important source

of soil acidity. When roots or soil organisms respire, they generate CO2. In wet soils, this CO2 goes into solution immediately, creating a weak acid, carbonic acid. This is also the reason why rainwater is naturally a bit acid: CO2 from the atmosphere dissolves into the raindrops. In tropical rainforests, where conditions favor massive amounts of respiration and there are often few clay particles to bind cations, this added CO2 can help make the soils quite acid and promotes the mobility of cations. The consequence is rapid uptake of nutrients by plants in undisturbed forests, and rapid loss of those nutrients when forests are cleared.

Horizons and Profiles

Soils are not homogenous; they contain characteristic layers, or horizons, that differ from one soil type to another. If you look at roadcuts or excavations dug for construction, you can often see the horizons quite clearly. The sequence of horizons that characterizes a soil is called the soil profile (Figure 4.4). The horizons are grouped under four categories: O, A, B, and C. Subcategories of these four categories are numbered according to their particular characteristics (Figure 4.5).

The O horizons consist of organic material formed above the mineral soil, derived from decaying plant materials, microbial matter, and the remains and waste products of animals. The A horizons—the surface layer of mineral soil— represent the region of maximum leaching, or eluviation. The uppermost A horizon, the A1, is often darker than the rest of the soil and may contain

Figure 4.3

Percentage base saturation in three soils: a clay loam (left), the same clay loam with agricultural lime (calcium carbonate) added to raise the pH (center), and a sandy loam with low cation exchange capacity (right). The CEC for each soil is indicated. CEC is measured in the SI units, centimoles of positive charge per kilogram of dry soil (cmolc/kg), indicating the numer of centimoles (1/100th of a mole) of positive charge adsorbed per unit mass of the soil; 1 mole of negative charge attracts 1 mole of positive charges whether the charges come from H+, K+, Ca2+ or Al3+. See a current soils text for further explanation. (After Buckman and Brady 1969.)

highly decayed organic matter. The B horizons, deeper in the soil, represent the region of maximum illuviation, or deposition of minerals and colloidal particles leached from elsewhere. Clays, iron, and aluminum oxides often accumulate in the B2 horizon. The C horizon is the undeveloped mineral material deep in the soil; it may or may not be the same as the material from which the soil develops. Below that may be bedrock, or just deep accumulations of mineral material deposited by wind, water, or glaciers.

No one soil has all of the horizons pictured in Figure 4.5. Some horizons may be much more distinct and better developed than others. Plowing and the activity of earthworms may obscure the distinction between the upper horizons. The soil profile may also not be fully developed, and some horizons may be absent or indistinct, if the soil is relatively young. Upper horizons may have been lost through erosion, leaving the deeper horizons exposed at the surface.

Erosion is a particular problem where forests have been clear-cut, on soils that have been long cultivated

Figure 4.4

Profile of a forest soil at the edge of the Adirondack Mountains in northern New York State (in the Spodosol soil order). This soil is a stony loam, forested with birch, hemlock, and spruce, and is quite acidic throughout the profile. There is a thick layer of organic material at the top of the profile, and a number of distinct horizons, each with particular properties and a characteristic appearance. Roots are shown reaching down to the top of the B3 horizon at about 45 cm deep, although the deeper roots of large trees would certainly penetrate more deeply into this soil. (After Buckman and Brady 1969.)

take thousands of years. The consequences of soil erosion for our ability to raise food crops and for the sustainability of natural ecosystems can be devastating.

Soils vary in depth, from thin layers of soil barely covering a rock substrate (for example, in many alpine areas) to very deep soils of close to 2 meters (for example, in some well-developed prairie soils). Soil depth has a great deal of influence on vegetation and plant growth: The greater the soil depth, the more favorable the soil for plant growth, all other things being equal. Deeper soils can hold more water and nutrients and can retain water for a longer period of time without rainfall, and they allow greater development of plant root systems.

Figure 4.5

An abstract, general soil profile showing the major horizons that might be present in particular soils. No one soil is likely to have all of the horizons shown, and particular soils may have greater development of subhorizons than what is shown. Only the upper part of the C horizon is considered to be part of the soil proper. The depth of soils varies tremendously with the location and nature of the soil, but to gain some perspective, readers might picture this illustrated profile as being about a meter deep to the top of the bedrock. (After Buckman and Brady 1969.)

Horizon

O1

O2

A1

A2

A3

B1

B2

B3

C

R

Properties

Organic; original forms recognized

Organic; original forms not recognized

Mineral mixed with humus; dark colored

Horizon of maximum eluviation of silicate clays; Fe, Al oxides, etc.

Transition to B; more like A than B Transition to A; more like B than A

Maximum illuviation of silicate clays; Fe, Al oxides, some organic matter

Transition to C; more like B than C

Zone of least weathering; accumulation of Ca, Mg carbonates; cementation

Bedrock

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DM looking for more up-to-date photo to replace this one.

Origins and Classification

Soils are formed by the action of weathering on parent material, the upper layers of the heterogeneous mass that is left over after the action of weathering and other forces on rocks. The physical and chemical actions of temperature, rainfall, and other climatic factors further grind up, move, and leach the parent material until it begins to develop into true soil. Parent material is originally derived from rocks, which are classified as igneous (of volcanic origin), sedimentary (from the deposition and recementation of material derived from other rocks), or metamorphic (changed by the action of great pressures and temperatures on igneous or sedimentary rocks deep underground).

The parent material can be broken up by the mechanical action of temperature changes, including repeated freezing and thawing, by the direct action of water, wind, and ice in abrading and eroding rock fragments, and by the actions of plants, fungi, and animals. Chemical reactions occurring in the disintegrating rock material accelerate the physical breakdown of the fragments and begin the process of decomposition and chemical alteration (Figure 4.7). Large and small fragments of rock and rock-derived debris are moved by ice (particularly by glaciers), wind, and water. They are then redeposited in the form of glacial till and outwash, loess (pronounced “luss”) and other aeolian (wind) deposits, alluvial (river and stream) deposits, and lake and marine sediments. These deposits may eventually form soil parent material.

The process of soil formation from parent material takes thousands of years. A young soil may be 10,000

Soils, Mineral Nutrition, and Belowground Interactions 69

Figure 4.6

During the 1930s, soil erosion in the midwestern United States led to massive dust storms that destroyed farms and even entire towns. What came to be known as the “Dust Bowl” displaced thousands of families, who left the region in search of a new livelihood. This shows the remains of a farm lot in South Dakota in 1936. (Photograph courtesy of the U.S. Department of Agriculture.)

years old; an old soil may be 100,000 years old or more. As a soil ages, the primary minerals that came from the parent material undergo chemical changes, and secondary minerals are formed. The structure of the soil develops, and then changes, as materials are leached, redeposited, and lost from the soil, as well as being physically and chemically altered. The mineral constituents of the soil shift, as minerals that are most stable at Earth’s surface gradually come to dominate. Young soils are found where the parent material is still present, as in areas that were glaciated relatively recently (such as much of the northern part of North America) or in areas of recent alluvial deposits (such as the bottomlands of the world’s great river systems). Old soils are most commonly found in the Tropics and Subtropics. Some of the world’s oldest soils are found, for instance, in parts of Africa.

Five major factors are responsible for determining the kinds of soils that develop in an area: climate, the nature of the parent material, the age of the soil, topography (which acts to alter the effects of climate),and living organisms. These factors do not operate in isolation, but rather interact in complex ways with one another. Vegetation, with its associated soil organisms, has an especially strong influence on soil development, but the nature of the vegetation that is present is also dependent on both soil and climate.

Soils are classified according to a comprehensive taxonomic system developed in the United States, although some soil scientists and ecologists still rely on earlier classification systems. The broadest category in the modern system is the soil order. There are ten soil orders worldwide (Table 4.1).

Figure 4.7

Development of soil structure from two different parent materials: (A) bedrock and (B) loess deposited by wind. These soils are also developing under two different climatic and vegetation types. Over time, organic matter accumulates in the upper horizon, depending on the type of vegetation present. In the lower horizons, clay, iron oxides, or CaCO3 are deposited and accumulate, and characteristic structures develop. The soil orders (ultisol and mollisol) are described in Table 4.1. (After Buckman and Brady 1969.)

The soil classification categories most commonly used by plant ecologists are the soil series and soil type. These are used to classify local soils. A soil series is usually named after some local geographic area or feature, and generally consists of a dominant type with several associated other types. Soil types are based upon topography, parent material, and the vegetation under which the soils were formed. Soil series occur at regional scales, and soil types characterize landscape-scale features. In the United States, soil series are mapped for almost every county, and detailed maps are available for most locations showing the soil series and type, and often providing extensive information on the characteristics of

local soils. Such data can be highly valuable for many kinds of ecological studies, particularly at the community and ecosystem levels. The availability of information on local soils in other countries varies widely.

Organic Matter and the Role of Organisms

So far we have emphasized the role of physical processes in soil development and character. However, we began this chapter by emphasizing that soils are the unique product of living things acting on the physical environment (and in turn being affected by that environment). What are some of the ways in which organisms affect soils?

Northeastern and northern midwestern United States and adjacent areas in Canada; northern Europe, Siberia; also some soils in southern South America

Large area from Baltic states to western Russia; southern half of Africa; eastern Brazil; England, France, central and parts of southern Europe; Michigan and Wisconsin to Pennsylvania and New York in the United States

One of the most widespread soil orders globally, and among the most important in terms of the size of the human population supported; also supports the greatest biodiversity of many groups of plants and animals. Covers much of Africa and South America

Sand hills of Nebraska; areas of northern and southern Africa; northern Quebec; parts of Siberia and Tibet; many areas that were covered by the most recent glaciation

All continents—north Africa, eastern China, western Siberia, Spain, central France, central Germany, northern South America, northwestern United States

Southwestern United States and northern Mexico; southern and central Australia; Bobi and Takla makan Deserts in China; Sahara Desert and also southwestern Africa; Pakistan; other arid regions

Large regions of India, Sudan, and eastern Australia; small areas in southeastern Texas and eastern Mississippi in the United States

Southeastern United States; northeastern Australia; Hawaii; southeastern Asia; southern Brazil

Not widespread, but may be locally important, as in the Everglades in southern Florida and in

many areas of northern Europe. The largest areas are in Canada southwest of Hudson and James Bay and in northwestern Canada into eastern Alaska

Note: These soil orders reflect the modern classification system currently used in the United States. Some of the major Great Soil Groups of the older classification system, with the soil orders with which they roughly correspond, are: podzols (spodosols and alfisols), chernozems and

brunizems (mollisols), latosols and lateritic soils (oxisols), lithosols, solonchak soils, and desert soils (aridisols), azonal soils (entisols), and wetland/bog soils (histisols).

Organic matter is the decaying and decomposed material in soil that comes from living things. It includes substances secreted by plants, microorganisms, and animals, products of excretion by animals, parts shed by animals and plants, and dead organisms and parts of organisms. Microscopic organisms, while tiny individually, collectively contribute a great deal of organic mat-

ter to the soil. The activities of various different kinds of microorganisms affect soil properties in many additional ways, including the recycling of mineral nutrients used by plants (see Chapter 15).

A property of soils that is of vital importance to plant life (and to animals, fungi, and the other organisms living in soils) is soil structure. Soil structure describes the