Chapters 2-4 Study Questions
CHAPTER 4
Smith, T. M., & Smith, R. L. (2015). Elements of Ecology (9th ed.). Boston, MA: Pearson.
4.1 Life on Land Imposes Unique Constraints
The transition from life in aquatic environments to life on land brought with it a variety of constraints. Perhaps the greatest constraint imposed by terrestrial environments is desiccation. Living cells, both plant and animal, contain about 75–95 percent water. Unless the air is saturated with moisture, water readily evaporates from the surfaces of cells via the process of diffusion (see Section 2.5). The water that is lost to the air must be replaced if the cell is to remain hydrated and continue to function. Maintaining this balance of water between organisms and their surrounding environment (referred to as an organism’s water balance) has been a major factor in the evolution of life on land. For example, in adapting to the terrestrial environment, plants have evolved extensively specialized cells for different functions. Aerial parts of most plants, such as stems and leaves, are coated with a waxy cuticle that prevents water loss. While it reduces water loss, the waxy surface also prevents gas exchange (carbon dioxide and oxygen) from occurring. As a result, terrestrial plants have evolved pores on the leaf surface (stomata) that allow gases to diffuse from the air into the interior of the leaf (see Chapter 6).
To stay hydrated, an organism must replace water that it has lost to the air. Terrestrial animals can acquire water by drinking and eating. For plants, however, the process is passive. Early in their evolution, land plants evolved vascular tissues consisting of cells joined into tubes that transport water and nutrients throughout the plant body. The topic of water balance and the array of characteristics that plants and animals have evolved to overcome the problems of water loss are discussed in more detail later (see Chapters 6 and 7).
The giant kelp (Macrocystis pyrifera) inhabits the waters off the coast of California. Anchored to the bottom sediments, these kelp plants can grow 100 feet or more toward the surface despite their lack of supportive tissues. These kelp plants are kept afloat through the buoyancy of gas-filled bladders attached to each blade, yet when the kelp plants are removed from the water, they collapse into a mass. (right) In contrast, a redwood tree (Sequoia sempervirens) of comparable height allocates more than 80 percent of its biomass to supportive and conductive tissues that help the tree resist gravitational forces.
Desiccation is not the only constraint imposed by the transition from water to land. Because air is less dense than water, it results in a much lower drag (frictional resistance) on the movement of organisms; but it greatly increases the constraint imposed by gravitational forces. The upward force of buoyancy resulting from the displacement of water helps organisms in aquatic environments overcome the constraints imposed by gravity (see Section 3.2). In contrast, the need to remain erect against gravitational force in terrestrial environments results in a significant investment in structural materials such as skeletons (for animals) or cellulose (for plants). The giant kelp (Macrocystis pyrifera) inhabiting the waters off the coast of California is an excellent example (Figure 4.1, left). It grows in dense stands called kelp forests. Anchored to the bottom sediments, these kelp (macroalgae) can grow 100 feet or more toward the surface. The kelp are kept afloat by gas-filled bladders attached to each blade; yet when the kelp plants are removed from the water, they collapse into a mass. Lacking supportive tissues strengthened by cellulose and lignin, the kelp cannot support its own weight under the forces of gravity. In contrast, a tree of equivalent height inhabiting the coastal forest of California (Figure 4.1, right) must allocate more than 80 percent of its total mass to supportive and conductive tissues in the trunk (bole), branches, leaves, and roots.
Another characteristic of terrestrial environments is their high degree of variability, both in time and space. Temperature variations on land (air) are much greater than in water. The high specific heat of water prevents wide daily and seasonal fluctuations in the temperature of aquatic habitats (see Section 3.2). In contrast, such fluctuations are a characteristic of air temperatures (see Chapter 2). Likewise, the timing and quantity of precipitation received at a location constrains the availability of water for terrestrial plants and animals as well as their ability to maintain water balance. These fluctuations in temperature and moisture have both a short-term effect on metabolic processes and a long-term influence on the evolution and distribution of terrestrial plants and animals (see Chapters 6 and 7). Ultimately, the geographic variation in climate governs the large-scale distribution of plants and therefore the nature of terrestrial ecosystems (see Chapter 23).
4.2 Plant Cover Influences the Vertical Distribution of Light
In contrast to aquatic environments, where the absorption of solar radiation by the water itself results in a distinct vertical gradient of light, the dominant factor influencing the vertical gradient of light in terrestrial environments is the absorption and reflection of solar radiation by plants. When walking into a forest in summer, you will observe a decrease in light (Figure 4.2a). You can observe much the same effect if you examine the lowest layer in grassland or an old field (Figure 4.2b). The quantity and quality (spectral composition) of light that does penetrate the canopy of vegetation to reach the ground varies with both the quantity and orientation of the leaves.
The amount of light at any depth in the canopy is affected by the number of leaves above. As we move down through the canopy, the number of leaves above increases; so the amount of light decreases. However, because leaves vary in size and shape, the number of leaves is not the best measure of quantity. The quantity of leaves, or foliage density, is generally expressed as the leaf area. Because most leaves are flat, the leaf area is the surface area of one or both sides of the leaf. When the leaves are not flat, the entire surface area is sometimes measured. To quantify the changes in light environment with increasing area of leaves, we need to define the area of leaves per unit ground area (m2 leaf area/m2 ground area). This measure is the leaf area index ([LAI]; Figure 4.3). A LAI of 3 indicates a quantity of 3 m2 of leaf area over each 1 m2 of ground area.
The greater the LAI above any surface, the lower the quantity of light reaching that surface. As you move from the top of the canopy to the ground in a forest, the cumulative leaf area and LAI increase. Correspondingly, light decreases. The general relationship between available light and LAI is described by Beer’s law (see Quantifying Ecology 4.1).
Absorption and reflection of light by the plant canopy. (a) A mixed conifer–deciduous forest reflects about 10 percent of the incident photosynthetically active radiation (PAR) from the upper canopy, and it absorbs most of the remaining PAR within the canopy. (b) A meadow reflects 20 percent of the PAR from the upper surface. The middle and lower regions, where the leaves are densest, absorb most of the rest. Only 2–5 percent of PAR reaches the ground.
The concept of leaf area index (LAI). (a) A tree with a crown 10 m wide projects a circle of the same size on the ground. (b) Foliage density (area of leaves) at various heights above the ground. (c) Contributions of layers in the crown to the leaf area index. (d) Calculation of leaf area index (LAI). The total leaf area is 315 m2. The projected ground area is 78.5 m2. The LAI is 4.
In addition to the quantity of light, the spectral composition (quality) of light varies through the plant canopy. Recall that the wavelengths of approximately 400 to 700 nm make up visible light (Section 2.1 and Figure 2.1). These wavelengths are also known as photosynthetically active radiation (PAR) because they include the wavelengths used by plants as a source of energy in photosynthesis (see Chapter 6). The transmittance of PAR is typically less than 10 percent, whereas the transmittance of far-red radiation (730 nm) is much greater. As a result, the ratio of red (660 nm) to far-red radiation (R/FR ratio) decreases through the canopy. This shift in the spectral quality of light affects the production of phytochrome (a pigment that allows a plant to perceive shading by other plants), thus influencing patterns of growth and allocation (see Chapter 6, Section 6.8).
Influence of leaf orientation (angle) on the interception of light energy. If a leaf that is perpendicular to the source of light (a) intercepts 1.0 unit of light energy, the same leaf at an angle of 60 degrees relative to the light source will intercept only 0.5 unit (b). The reduction in intercepted light energy is a result of the angled leaf projecting a smaller surface area relative to the light source.
Besides the quantity of leaves, the orientation of leaves on the plant influences the attenuation of light through the canopy. The angle at which a leaf is oriented relative to the Sun changes the amount of light it absorbs. If a leaf that is perpendicular to the Sun absorbs 1.0 unit of light energy (per unit leaf area/time), the same leaf displayed at a 60-degree angle to the Sun will absorb only 0.5 units. The reason is that the same leaf area represents only half the projected surface area and therefore intercepts only half as much light energy (Figure 4.4). Thus, leaf angle influences the vertical distribution of light through the canopy as well as the total amount of light absorbed and reflected. The sun angle varies, however, both geographically (see Section 2.1) and through time at a given location (over the course of the day and seasonally). Consequently, different leaf angles are more effective at intercepting light in different locations and at different times. For example, in high-latitude environments, where sunlight angles are low, canopies having leaves that are displayed at an angle will absorb light more effectively (see Figure 2.5). Leaves that are displayed at an angle rather than perpendicular to the Sun are also typical of arid tropical environments. In these hot and dry environments, angled leaves reduce light interception during midday, when temperatures and demand for water are at their highest.
Although light decreases downward through the plant canopy, some direct sunlight does penetrate openings in the crown and reaches the ground as sunflecks. Sunflecks can account for 70–80 percent of solar energy reaching the ground in forest environments (Figure 4.5).
Quantifying Ecology 4.1 Beer’s Law and the Attenuation of Light
Due to the absorption and reflection of light by leaves, there is a distinct vertical gradient of light availability from the top of a plant canopy to the ground. The greater the surface area of leaves, the less light will penetrate the canopy and reach the ground. The vertical reduction, or attenuation, of light through a stand of plants can be estimated using Beer’s law, which describes the attenuation of light through a homogeneous medium. The medium in this case is the canopy of leaves. Beer’s law can be applied to the problem of light attenuation through a plant canopy using the following relationship:
The subscript i refers to the vertical height of the canopy. For example, if i were in units of meters, a value of i = 5 refers to a height of 5 m above the ground. The value e is the natural logarithm (2.718). The light extinction coefficient, k, represents the quantity of light attenuated per unit of leaf area index (LAI) and is a measure of the degree to which leaves absorb and reflect light. The extinction coefficient will vary as a function of leaf angle (see Figure 4.4) and the optical properties of the leaves. Although the value of ALi is expressed as a proportion of the light reaching the top of the canopy, the quantity of light at any level can be calculated by multiplying this value by the actual quantity of light (or photosynthetically active radiation) reaching the top of the canopy (units of μmol/m2/s).
For the example presented in Figure 4.3, we can construct a curve describing the available light at any height in the canopy. In Figure 1, the light extinction coefficient has a value of k = 0.6 as an average value for a temperate deciduous forest. We label vertical positions from the top of the canopy to ground level on the curve. Knowing the amount of leaves (LAI) above any position in the canopy (i), we can use the equation to calculate the amount of light there.
The availability of light at any point in the canopy will directly influence the levels of photosynthesis (see Figure 6.2). The light levels and rates of light-limited photosynthesis for each of the vertical canopy positions are shown in the curve in Figure 2. Light levels are expressed as a proportion of values for fully exposed leaves at the top of the canopy (1500 μmol/m2/s). As one moves from the top of the canopy downward, the amount of light reaching the leaves and the corresponding rate of photosynthesis decline.
Beer’s law can also be used to describe the vertical attenuation of light in aquatic environments, but applying the light extinction coefficient (k) is more complex. The reduction of light with water depth is a function of various factors: (1) attenuation by the water itself (see Section 3.3, Figure 3.7); (2) attenuation by phytoplankton (microscopic plants suspended in water), typically expressed as the concentration of chlorophyll (the light-harvesting pigment of plants) per volume of water (see Section 6.1); (3) attenuation by dissolved substances; and (4) attenuation by suspended particulates. Each of these factors has an associated light extinction coefficient, and the overall light extinction coefficient (kT) is the sum of the individual coefficients:
Whereas the light extinction coefficient for leaf area expresses the attenuation of light per unit of LAI, these values of k are expressed as the attenuation of light per unit of water depth (such as centimeter, meter, inches, or feet). Beer’s law can then be used to estimate the quantity of light reaching any depth (z) by using the following equation:
ALZ=e−kTZALZ= e−kTZ
If the ecosystem supports submerged vegetation, such as kelp (see Figure 4.1), seagrass, or other plants that are rooted in the bottom sediments, the preceding equation can be used to calculate the available light at the top of the canopy. The equation describing the attenuation of light as a function of LAI can then be applied (combined) to calculate the further attenuation from the top of the plant canopy to the sediment surface.
If we assume that the value of k used to calculate the vertical profile of light in Figure 1 (k = 0.6) is for a plant canopy where the leaves are positioned horizontally (parallel to the forest floor), how would the value of k differ (higher or lower) for a forest where the leaves were oriented at a 60-degree angle? (See the example in Figure 4.4.)
In shallow-water ecosystems, storms and high wind can result in bottom sediments (particulates) being suspended in the water for some time before once again settling to the bottom. How would this situation affect the value of k T and the attenuation of light in the water profile?
In many environments, seasonal changes strongly influence leaf area. For example, in the temperate regions of the world, many forest tree species are deciduous, shedding their leaves during the winter months. In these cases, the amount of light that penetrates a forest canopy varies with the season (Figure 4.6). In early spring in temperate regions, when leaves are just expanding, 20–50 percent of the incoming light may reach the forest floor. In other regions characterized by distinct wet and dry seasons, a similar pattern of increased light availability at the ground level occurs during the dry season (see Chapter 2).
4.3 Soil Is the Foundation upon which All Terrestrial Life Depends
Soil is the medium for plant growth; the principal factor controlling the fate of water in terrestrial environments; nature’s recycling system, which breaks down the waste products of plants and animals and transforms them into their basic elements; and a habitat to a diversity of animal life, from small mammals to countless forms of microbial life (see Chapter 21).
As familiar as it is, soil is hard to define. One definition says that soil is a natural product formed and synthesized by the weathering of rocks and the action of living organisms. Another states that soil is a collection of natural bodies of earth, composed of mineral and organic matter and capable of supporting plant growth. Indeed, one eminent soil scientist, Hans Jenny—a pioneer of modern soil studies—will not give an exact definition of soil. In his book The Soil Resource, he writes:
Popularly, soil is the stratum below the vegetation and above hard rock, but questions come quickly to mind. Many soils are bare of plants, temporarily or permanently, or they may be at the bottom of a pond growing cattails. Soil may be shallow or deep, but how deep? Soil may be stony, but surveyors (soil) exclude the larger stones. Most analyses pertain to fine earth only. Some pretend that soil in a flowerpot is not a soil, but soil material. It is embarrassing not to be able to agree on what soil is. In this, soil scientists are not alone. Biologists cannot agree on a definition of life and philosophers on philosophy.
Of one fact we are sure. Soil is not just an abiotic environment for plants. It is teeming with life—billions of minute and not so minute animals, bacteria, and fungi. The interaction between the biotic and the abiotic makes the soil a living system.
Soil scientists recognize soil as a three-dimensional unit, or body, having length, width, and depth. In most places on Earth’s surface, exposed rock has crumbled and broken down to produce a layer of unconsolidated debris overlaying the hard, unweathered rock. This unconsolidated layer, called the regolith, varies in depth from virtually nonexistent to tens of meters. This interface between rock and the air, water, and living organisms that characterizes the surface environment is where soil is formed.
4.4 The Formation of Soil Begins with Weathering
Soil formation begins with the weathering of rocks and their minerals. Weathering includes the mechanical destruction of rock materials into smaller particles as well as their chemical modification. Mechanical weathering results from the interaction of several forces. When exposed to the combined action of water, wind, and temperature, rock surfaces flake and peel away. Water seeps into crevices, freezes, expands, and cracks the rock into smaller pieces. Wind-borne particles, such as dust and sand, wear away at the rock surface. Growing roots of trees split rock apart.
Without appreciably influencing their composition, mechanical weathering breaks down rock and minerals into smaller particles. Simultaneously, these particles are chemically altered and broken down through chemical weathering. The presence of water, oxygen, and acids resulting from the activities of soil organisms and the continual addition of organic matter (dead plant and animal tissues) enhance the chemical weathering process. Rainwater falling on and filtering through this organic matter and mineral soil sets up a chain of chemical reactions that transform the composition of the original rocks and minerals.
4.5 Soil Formation Involves Five Interrelated Factors
Five interdependent factors are important in soil formation: parent material, climate, biotic factors, topography, and time. Parent material is the material from which soil develops. The original parent material could originate from the underlying bedrock; from glacial deposits (till); from sand and silt carried by the wind (eolian); from gravity moving material down a slope (colluvium); and from sediments carried by flowing water (fluvial), including water in floodplains. The physical character and chemical composition of the parent material are important in determining soil properties, especially during the early stages of development.
Biotic factors—plants, animals, bacteria, and fungi—all contribute to soil formation. Plant roots can function to break up parent material, enhancing the process of weathering, as well as stabilizing the soil surface and reducing erosion. Plant roots pump nutrients up from soil depths and add them to the surface. In doing so, plants recapture minerals carried deep into the soil by weathering processes. Through photosynthesis, plants capture the Sun’s energy and transfer some of this energy to the soil in the form of organic carbon. On the soil surface, microorganisms break down the remains of dead plants and animals that eventually become organic matter incorporated into the soil (see Chapter 21). Climate influences soil development both directly and indirectly. Temperature, precipitation, and winds directly influence the physical and chemical reactions responsible for breaking down parent material and the subsequent leaching (movement of solutes through the soil) and movement of weathered materials. Water is essential for the process of chemical weathering, and the greater the depth of water percolation, the greater the depth of weathering and soil development. Temperature controls the rates of biochemical reactions, affecting the balance between the accumulation and breakdown of organic materials. Consequently, under conditions of warm temperatures and abundant water, the processes of weathering, leaching, and plant growth (input of organic matter) are maximized. In contrast, under cold, dry conditions, the influence of these processes is much more modest. Indirectly, climate influences a region’s plant and animal life, both of which are important in soil development.
Topography, the contour of the land, can affect how climate influences the weathering process. More water runs off and less enters the soil on steep slopes than on level land; whereas water draining from slopes enters the soil on low and flat land. Steep slopes are also subject to soil erosion and soil creep—the downslope movement of soil material that accumulates on lower slopes and lowlands.
Time is a crucial element in soil formation: all of the factors just listed assert themselves over time. The weathering of rock material; the accumulation, decomposition, and mineralization of organic material; the loss of minerals from the upper surface; and the downward movement of materials through the soil all require considerable time. Forming well-developed soils may require 2000 to 20,000 year
4.6 Soils Have Certain Distinguishing Physical Characteristics
Soils are distinguished by differences in their physical and chemical properties. Physical properties include color, texture, structure, moisture, and depth. All may be highly variable from one soil to another.
Color is one of the most easily defined and useful characteristics of soil. It has little direct influence on the function of a soil but can be used to relate chemical and physical properties. Organic matter (particularly humus) makes soil dark or black. Other colors can indicate the chemical composition of the rocks and minerals from which the soil was formed. Oxides of iron give a color to the soil ranging from yellowish-brown to red, whereas manganese oxides give the soil a purplish to black color. Quartz, kaolin, gypsum, and carbonates of calcium and magnesium give whitish and grayish colors to the soil. Blotches of various shades of yellowish-brown and gray indicate poorly drained soils or soils saturated by water. Soils are classified by color using standardized color charts (i.e., Munsell soil color charts).
Soil texture is the proportion of different-sized soil particles. Texture is partly inherited from parent material and partly a result of the soil-forming process. Particles are classified on the basis of size into gravel, sand, silt, and clay. Gravel consists of particles larger than 2.0 mm, but they are not part of the fine fraction of soil. Soils are classified based on texture by defining the proportion of sand, silt, and clay.
Sand ranges from 0.05 to 2.0 mm, is easy to see, and feels gritty. Silt consists of particles from 0.002 to 0.05 mm in diameter that can scarcely be seen by the naked eye; it feels and looks like flour. Clay particles are less than 0.002 mm and are too small to be seen under an ordinary microscope. Clay controls the most important properties of soils, including its water-holding capacity (see Section 4.8) and the exchange of ions between soil particles and soil solution (see Section 4.9). A soil’s texture is the percentage (by weight) of sand, silt, and clay. Based on proportions of these components, soils are divided into texture classes (Figure 4.7).
Interpreting Ecological Data
Q1. What is the texture classification for a soil with 60 percent silt, 35 percent clay, and 5 percent sand?
Q2. What is the texture classification for a soil with 60 percent clay and 40 percent silt?
Soil texture affects pore space in the soil, which plays a major role in the movement of air and water in the soil and the penetration by roots. In an ideal soil, particles make up 50 percent of the soil’s total volume; the other 50 percent is pore space. Pore space includes spaces within and between soil particles, as well as old root channels and animal burrows. Coarse-textured soils have large pore spaces that favor rapid water infiltration, percolation, and drainage. To a point, the finer the texture, the smaller the pores, and the greater the availability of active surface for water adhesion and chemical activity. Very fine-textured or heavy soils, such as clays, easily become compacted if plowed, stirred, or walked on. They are poorly aerated and difficult for roots to penetrate.
Soil depth varies across the landscape, depending on slope, weathering, parent materials, and vegetation. In grasslands, much of the organic matter added to the soil is from the deep, fibrous root systems of the grass plants. By contrast, tree leaves falling on the forest floor are the principal source of organic matter in forests. As a result, soils developed under native grassland tend to be several meters deep, and soils developed under forests are shallow. On level ground at the bottom of slopes and on alluvial plains, soils tend to be deep. Soils on ridgetops and steep slopes tend to be shallow, with bedrock close to the surface.
4.7 The Soil Body Has Horizontal Layers or Horizons
Initially, soil develops from undifferentiated parent material. Over time, changes occur from the surface down, through the accumulation of organic matter near the surface and the downward movement of material. These changes result in the formation of horizontal layers that are differentiated by physical, chemical, and biological characteristics. Collectively, a sequence of horizontal layers constitutes a soil profile. This pattern of horizontal layering, or horizons, is easily visible where a recent cut has been made along a road bank or during excavation for a building site (Figure 4.8).
The pattern of horizontal layering or soil horizons is easily visible where a recent cut has been made along a road bank. This soil is relatively shallow, with the parent material close to the surface.
The simplest general representation of a soil profile consists of four horizons: O, A, B, and C (Figure 4.9). The surface layer is the O horizon, or organic layer. This horizon is dominated by organic material, consisting of partially decomposed plant materials such as leaves, needles, twigs, mosses, and lichens. This horizon is often subdivided into a surface layer composed of undecomposed leaves and twigs (Oi), a middle layer composed of partially decomposed plant tissues (Oe), and a bottom layer consisting of dark brown to black, homogeneous organic material or the humus layer (Oa). This pattern of layering is easily seen by carefully scraping away the surface organic material on the forest floor. In temperate regions, the organic layer is thickest in the fall, when new leaf litter accumulates on the surface. It is thinnest in the summer after decomposition has taken place.
A generalized soil profile. Over time, changes occur from the surface down, through the accumulation of organic matter near the surface and the downward movement of material. These changes result in the formation of horizontal layers, or horizons.
Below the organic layer is the A horizon, often referred to as the topsoil. This is the first of the layers that are largely composed of mineral soil derived from the parent materials. In this horizon, organic matter (humus) leached from above accumulates in the mineral soil. The accumulation of organic matter typically gives this horizon a darker color, distinguishing it from lower soil layers. Downward movement of water through this layer also results in the loss of minerals and finer soil particles, such as clay, to lower portions of the profile—sometimes giving rise to an E horizon, a zone or layer of maximum leaching, or eluviation (from Latin ex or e, “out,” and lavere, “to wash”) of minerals and finer soil particles to lower portions of the profile. Such E horizons are quite common in soils developed under forests, but because of lower precipitation they rarely occur in soils developed under grasslands.
Below the A (or E) horizon is the B horizon, also called the subsoil. Containing less organic matter than the A horizon, the B horizon shows accumulations of mineral particles such as clay and salts from the leaching from the topsoil. This process is called illuviation (from the Latin il, “in,” and lavere, “to wash”). The B horizon usually has a denser structure than the A horizon, making it more difficult for plants to extend their roots downward. B horizons are distinguished on the basis of color, structure, and the kind of material that has accumulated as a result of leaching from the horizons above.
The C horizon is the unconsolidated material that lies under the subsoil and is generally made of original material from which the soil developed. Because it is below the zones of greatest biological activity and weathering and has not been sufficiently altered by the soil-forming processes, it typically retains much of the characteristics of the parent materials from which it was formed. Below the C horizon lies the bedrock.
4.8 Moisture-Holding Capacity Is an Essential Feature of Soils
If you dig into the surface layer of a soil after a soaking rain, you should discover a sharp transition between wet surface soil and the dry soil below. As rain falls on the surface, it moves into the soil by infiltration. Water moves by gravity into the open pore spaces in the soil, and the size of the soil particles and their spacing determine how much water can flow in. Wide pore spacing at the soil surface increases the rate of water infiltration; so coarse soils have a higher infiltration rate than fine soils do.
If there is more water than the pore space can hold, we say that the soil is saturated, and excess water drains freely from the soil. If water fills all the pore spaces and is held there by internal capillary forces, the soil is at field capacity (physically defined as the water content at –0.33 bar suction pressure, or .0033 MPa). Field capacity is generally expressed as the percentage of the weight or volume of soil occupied by water when saturated compared to the oven-dried weight of the soil at a standard temperature. The amount of water a soil holds at field capacity varies with the soil’s texture—the proportion of sand, silt, and clay. Coarse, sandy soil has large pores; water drains through it quickly. Clay soils have small pores and hold considerably more water. Water held between soil particles by capillary forces is capillary water.
As plants and evaporation from the soil surface extract capillary water, the amount of water in the soil declines. When the moisture level decreases to a point where plants can no longer extract water, the soil has reached the wilting point (physically defined as the water content at –15 bar suction pressure, or –1.5 MPa). The amount of water retained by the soil between field capacity and wilting point (or the difference between field capacity and wilting point) is the available water capacity (AWC), as shown in Figure 4.10. The AWC provides an estimate of the water available for uptake by plants. Although water still remains in the soil—filling up to 25 percent of the pore spaces—soil particles hold it tightly, making it difficult to extract.
Water content of three different soils at wilting point (WP), field capacity (FC), and saturation. The three soils differ in texture from coarse-textured sand to fine-textured silty clay loam (see soil texture chart of Figure 4.7). Available water capacity (AWC) is defined as the difference between FC and WP. Both FC and WP increase from coarse- to fine-textured soils, and the highest AWC is in the intermediate-textured soils.
Interpreting Ecological Data
Q1. Although fine-textured soils (silty clay loam) have a greater AWC, for this value to be achieved, the soil must be at or above FC. In arid regions, low and infrequent precipitation may keep soil water content below FC for most of the growing season. If the measured value of soil water content at a site is 10 g/cm3, which soil texture (sand, silt, or clay) represented in Figure 4.10 would have the greatest soil water available for uptake by plants?
Q2. What if the value of soil water was 35 g/cm3?
Both the field capacity and wilting point of a soil are heavily influenced by soil texture. Particle size of the soil directly influences the pore space and surface area onto which water adheres. Sand has 30–40 percent of its volume in pore space, whereas clays and loams (see soil texture chart in Figure 4.7) range from 40 to 60 percent. As a result, fine-textured soils have a higher field capacity than sandy soils, but the increased surface area results in a higher value of the wilting point as well (see Figure 4.10). Conversely, coarse-textured soils (sands) have a low field capacity and a low wilting point. Thus, AWC is highest in intermediate clay loam soils.
The topographic position of a soil affects the movement of water both on and in the soil. Water tends to drain downslope, leaving soils on higher slopes and ridgetops relatively dry and creating a moisture gradient from ridgetops to streams.
4.9 Ion Exchange Capacity Is Important to Soil Fertility
Chemicals within the soil dissolve into the soil water to form a solution (see Section 3.5). Referred to as exchangeable nutrients, these chemical nutrients in solution are the most readily available for uptake and use by plants (see Chapter 6). They are held in soil by the simple attraction of oppositely charged particles and are constantly interchanging with the soil solution.
As described previously, an ion is a charged particle. Ions carrying a positive charge are cations, and ions carrying a negative charge are anions. Chemical elements and compounds exist in the soil solution both as cations, such as calcium (Ca2+), magnesium (Mg2+), and ammonium (NH4+), and as anions, such as nitrate (NO3−) and sulfate (SO42−). The ability of these ions in soil solution to bind to the surface of soil particles depends on the number of negatively or positively charged sites within the soil. The total number of charged sites on soil particles within a volume of soil is called the ion exchange capacity. In most soils of the temperate zone, cation exchange predominates over anion exchange because of the prevalence of negatively charged particles in the soil, referred to as colloids. The total number of negatively charged sites, located on the leading edges of clay Particles and soil organic matter (humus particles), is called the cation exchange capacity (CEC). These negative charges enable a soil to prevent the leaching of its positively charged nutrient cations. Because in most soils there are far fewer positively charged than negatively charged sites, anions such as nitrate (NO3−) and phosphate (PO34−) are not retained on exchange sites in soils but tend to leach away quickly if not taken up by plants. The CEC is a basic measure of soil quality and increases with higher clay and organic matter content.
Cations occupying the negatively charged particles in the soil are in a state of dynamic equilibrium with similar cations in the soil solution (Figure 4.11). Cations in soil solution are continuously being replaced by or exchanged with cations on the clay and humus particles. The relative abundance of different ions on exchange sites is a function of their concentration in the soil solution and the relative affinity of each ion for the sites. In general, the physically smaller the ion and the greater its positive charge, the more tightly it is held. The lyotropic series places the major cations in order of their strength of bonding to the cation exchange sites in the soil:
The process of cation exchange in soils. Cations occupying the negatively charged particles in the soil are in a state of dynamic equilibrium with similar cations in the soil solution. Cations in soil solution are continuously being replaced by or exchanged with cations on clay and humus particles. Cations in the soil solution are also taken up by plants and leached to ground and surface waters.
AI3+>H+>Ca2+>Mg2+>k+=NH+4>Na+AI3+ >H+ >Ca2+>Mg2+>k+ =NH4+>Na+
However, higher concentrations in the soil solution can overcome these differences in affinity.
Hydrogen ions added by rainwater, by acids from organic matter, and by metabolic acids from roots and microorganisms increase the concentration of hydrogen ions in the soil solution and displace other cations, such as Ca2+, on the soil exchange sites. As more and more hydrogen ions replace other cations, the soil becomes increasingly acidic (see Section 3.7). Acidity is one of the most familiar of all chemical conditions in the soil. Typically, soils range from pH 3 (extremely acid) to pH 9 (strongly alkaline). Soils of more than pH 7 (neutral) are considered basic, and those of pH 5.6 or less are acid. As soil acidity increases, the proportion of exchangeable Al3+ increases, and Ca2+, Na+, and other cations decrease. High aluminum (Al3+) concentrations in soil solution can be toxic to plants. Aluminum toxicity damages the root system first, making the roots short, thick, and stubby. The result is reduced nutrient uptake.
4.10 Basic Soil Formation Processes Produce Different Soils
Broad regional differences in geology, climate, and vegetation give rise to characteristically different soils. The broadest level of soil classification is the order. Each order has distinctive features, summarized in Figure 4.12, and its own distribution, mapped in Figure 4.13. Although a wide variety of processes are involved in soil formation (pedogenesis), soil scientists recognize five main soil-forming processes that give rise to these different classes of soils. These processes are laterization, calcification, salinization, podzolization, and gleization.
Laterization is a process common to soils found in humid environments in the tropical and subtropical regions. The hot, rainy conditions cause rapid weathering of rocks and minerals. Movements of large amounts of water through the soil cause heavy leaching, and most of the compounds and nutrients made available by the weathering process are transported out of the soil profile if not taken up by plants. The two exceptions to this process are compounds of iron and aluminum. Iron oxides give tropical soils their unique reddish coloring (see Ultisol profile in Figure 4.12). Heavy leaching also causes these soils to be acidic because of the loss of other cations (other than H1).
(Adapted from USGS, Soil Conservation Service.)
Calcification occurs when evaporation and water uptake by plants exceed precipitation. The net result is an upward movement of dissolved alkaline salts, typically calcium carbonate (CaCO3), from the groundwater. At the same time, the infiltration of water from the surface causes a downward movement of the salts. The net result is the deposition and buildup of these deposits in the B horizon (subsoil). In some cases, these deposits can form a hard layer called caliche (Figure 4.14 top).
Salinization is a process that functions similar to calcification, only in much drier climates. It differs from calcification in that the salt deposits occur at or near the soil surface (Figure 4.14 bottom). Saline soils are common in deserts but may also occur in coastal regions as a result of sea spray. Salinization is also a growing problem in agricultural areas where irrigation is practiced.
Podzolization occurs in cool, moist climates of the midlatitude regions where coniferous vegetation (e.g., pine forests) dominates. The organic matter of coniferous vegetation creates strongly acidic conditions. The acidic soil solution enhances the process of leaching, causing the removal of cations and compounds of iron and aluminum from the A horizon (topsoil). This process creates a sublayer in the A horizon that is composed of white- to gray-colored sand (see Spodosol profile in Figure 4.12).
Gleization occurs in regions with high rainfall or low-lying areas associated with poor drainage (waterlogged). The constantly wet conditions slow the breakdown of organic matter by decomposers (bacteria and fungi), allowing the matter to accumulate in upper layers of the soil. The accumulated organic matter releases organic acids that react with iron in the soil, giving the soil a black to bluish-gray color (see Gelisol profile in Figure 4.12 as an example of soil formed through the process of gleization).
These five processes represent the integration of climate and edaphic (relating to the soil) factors on the formation of soils, giving rise to the geographic diversity of soils that influence the distribution, abundance, and productivity of terrestrial ecosystems. (We will explore these topics further in Chapters 20, 21, and 23.)
(top) In arid regions, salinization occurs when salts (the white crust at the center of the photo) accumulate near the soil surface because of surface evaporation. (bottom) Calcification occurs when calcium carbonates precipitate out from water moving downward through the soil or from capillary water moving upward from below. The result is an accumulation of calcium in the B horizon (seen as the white soil layer in the photo).
Ecological Issues & Applications Soil Erosion Is a Threat to Agricultural Sustainability
In a report released in 1909, the U.S. Bureau of Soils stated “The soil is the one indestructible, immutable asset that the nation possesses. It is the one resource that cannot be exhausted; that cannot be used up.” Yet less than three decades later, the loss of soil resources would be at the center of one of the worst environmental disasters in U.S. history—the Dust Bowl; a disaster that would have profound economic, social, and environmental costs.
Between 1909 and 1929 farmers had tilled some 13 million hectares of land in the Great Plains. In doing so they destroyed the sod—the grass-covered surface soil held together by the dense mat of fibrous roots. Once this protective cover of the native grassland was destroyed, the severe drought conditions and high winds during the period of the 1930s resulted in an increased susceptibility of the topsoil to wind erosion. As a result, dust storms raged nearly everywhere across the Great Plains of North America; but the most severely affected areas were in the Oklahoma and Texas panhandles, western Kansas, eastern Colorado, and northeastern New Mexico—a region that would become known as the Dust Bowl (Figure 4.15a). The most severe dust storms occurred between 1935 and 1938, although they would continue through 1941. It was estimated that 300 million tons of soil were removed from the region in May 1934 and spread over large portions of the eastern United States. By 1935 an additional 850 million tons of topsoil were removed by wind erosion. It is estimated that by 1935 wind erosion had damaged 66 million hectares across 80 percent of the High Plains. By 1938 it was estimated that 12.5 inches of topsoil had been lost over an area of 4 million hectares and 6.5 cm had been lost over another 5.5 million hectares.
The storms generated by this environmental disaster darkened cities, buried homes and farm equipment, killed livestock, and represented a serious health risk (Figure 4.15b and c). Overall, the Dust Bowl rendered millions of acres of farmland virtually useless, left roughly half a million Americans homeless, and forced hundreds of thousands of people off the land. It also resulted in the most intense period of internal migration in U.S. history. Between 1932 and 1940, it is estimated that 2.5 million people abandoned the plains for other regions of the country.
In response to the environmental disaster of the Dust Bowl, U.S. president, Franklin Delano Roosevelt, established the Soil Erosion Service (later the Soil Conservation Service, and now the Natural Resources Conservation Service), which marked the first major federal commitment to the preservation of natural resources in private hands. Even more significantly, in 1935, the Prairie States Forestry Project was established. Under this federal project, nearly 220 million trees were planted, creating more than 18,000 miles of windbreaks on some 30,000 farms, which formed a “shelter belt” from the Texas Panhandle to the Canadian border.
Although the end of the drought, together with soil conservation efforts following the Dust Bowl, abated the dramatic dust storms that blackened the skies over North America, the problem of soil erosion on agricultural lands remains a serious environmental issue. Approximately 50 percent of Earth’s land surface is devoted to agriculture, with about one-third planted in crops and two-thirds used for grazing. Of these two areas, cropland is more susceptible to erosion because the vegetation is most often removed and the soil tilled (plowed) before crops are planted. This functions to destabilize the soil surface, increasing rates of erosion resulting from both wind and water (Figure 4.16a). In addition, croplands are often left without vegetation cover between plantings (exposing the bare soil surface to erosion). According to David Pimentel of Cornell University, one of the leading experts in the study of agricultural ecology, currently about 80 percent of the world’s agricultural land suffers moderate to severe soil erosion. Worldwide, erosion on cropland averages about 30 tons per hectare per year and ranges from 0.5 to 400 tons per hectare per year. As a result of soil erosion, during the past four decades about 30 percent of the world’s arable land has become unproductive, much of which has been abandoned for agricultural use. Each year an estimated 10 million hectares of cropland worldwide are abandoned because of lack of productivity caused by soil erosion.
Rates of soil erosion on agricultural lands are influenced by a variety of factors. Topography of the landscape, patterns of rainfall and wind, and exposure all combine to influence the susceptibility of the soil surface to erosion. Soil structure influences the ease with which soils can be eroded. Soils with medium-to-fine texture (see Section 4.6) and low organic matter content are most easily eroded. Typically these soils have low water infiltration rates and are therefore susceptible to high rates of erosion by water and displacement by wind. Plant cover, both living and dead, greatly reduces rates of erosion by protecting the soil surface from exposure to agents of erosion.
Current estimates suggest that the degradation of agricultural lands alone will depress world food production by approximately 30 percent over the next 50 years, while during that same period the world population is predicted to exceed 9 billion (United Nations medium scenario; see Chapter 11, Ecological Issues & Applications). These forecasts point to the need to develop soil conservation techniques known to dramatically reduce soil erosion. For example, commercial corn production in the United States, which uses a practice of continuous crop production with annual plowing and removal of all plant materials at harvest, results in an average soil erosion rate of 44 tons per hectare per year. By using a practice of crop rotation in which a series of dissimilar/different types of crops are planted in the same area in sequential seasons (e.g., corn, wheat, and hay) erosion rates have been shown to decline to as little as 3 tons per hectare per year. No-till techniques, in which crops are planted directly in the soil without tilling or plowing the ground (Figure 4.16b), reduce average rates of erosion to 0.14 tons per hectare per year in corn fields. Similar reductions in rates of erosion have been measured with contour planting (plowing and planting row crops on a contour rather than up and down hill; Figure 4.16c) and the use of grass strips between crop rows. What all of these techniques share in common is that they serve to protect the soil surface from direct exposure to wind and rain.
Summary
Life on Land 4.1
Maintaining the balance of water between organisms and their surrounding environment has been a major influence on the evolution of life on land. The need to remain erect against the force of gravity in terrestrial environments results in a significant investment in structural materials. Variations in temperature and precipitation have both a short-term effect on metabolic processes and a long-term influence on the evolution and distribution of terrestrial plants and animals. The result is a distinct pattern of terrestrial ecosystems across geographic gradients of temperature and precipitation.
Light 4.2
Light passing through a canopy of vegetation becomes attenuated. The density and orientation of leaves in a plant canopy influence the amount of light reaching the ground. Foliage density is expressed as leaf area index (LAI), the area of leaves per unit of ground area. The amount of light reaching the ground in terrestrial vegetation varies with the season. In forests, only about 1–5 percent of light striking the canopy reaches the ground. Sunflecks on the forest floor enable plants to endure shaded conditions.
Soil Defined 4.3
Soil is a natural product of unconsolidated mineral and organic matter on Earth’s surface. It is the medium for plant growth; the principal factor controlling the fate of water in terrestrial environments; nature’s recycling system, which breaks down the waste products of plants and animals and transforms them into their basic elements; and a habitat to a diversity of animal life.
Weathering 4.4
Soil formation begins with the weathering of rock and minerals. In mechanical weathering, water, wind, temperature, and plants break down rock. In chemical weathering, the activity of soil organisms, the acids they produce, and rainwater break down primary minerals.
Soil Formation 4.5
Soil results from the interaction of five factors: parent material, climate, biotic factors, topography, and time. Parent material provides the substrate from which soil develops. Climate shapes soil development through temperature, precipitation, and its influence on vegetation and animal life. Biotic factors—vegetation, animals, bacteria, and fungi—add organic matter and mix it with mineral matter. Topography influences the amount of water entering the soil and the rates of erosion. Time is required to fully develop distinctive soils.
Distinguishing Characteristics 4.6
Soils differ in the physical properties of color, texture, and depth. Although color has little direct influence on soil function, it can be used to relate chemical and physical properties. Soil texture is the proportion of different-sized soil particles—sand, silt, and clay. A soil’s texture is largely determined by the parent material but is also influenced by the soil-forming process. Soil depth varies across the landscape, depending on slope, weathering, parent materials, and vegetation.
Soil Horizons 4.7
Soils develop in layers called horizons. Four horizons are commonly recognized, although not all of them are necessarily present in any one soil: the O or organic layer; the A (sometimes E) horizon, or topsoil, characterized by accumulation of organic matter; the B horizon, or subsoil, in which mineral materials accumulate; and the C horizon, the unconsolidated material underlying the subsoil and extending downward to the bedrock.
Moisture-Holding Capacity 4.8
The amount of water a soil can hold is one of its important characteristics. When water fills all pore spaces, the soil is saturated. When a soil holds the maximum amount of water it can retain, it is at field capacity. Water held between soil particles by capillary forces is capillary water. When the moisture level is at a point where plants cannot extract water, the soil has reached wilting point. The amount of water retained between field capacity and wilting point is the available water capacity. The available water capacity of a soil is a function of its texture.
Ion Exchange 4.9
Soil particles, particularly clay particles and organic matter, are important to nutrient availability and the cation exchange capacity of the soil—the number of negatively charged sites on soil particles that can attract positively charged ions. Cations occupying the negatively charged particles in the soil are in a state of dynamic equilibrium with similar cations in the soil solution. Percent base saturation is the percentage of sites occupied by ions other than hydrogen.
Soil Formation Processes Form Different Soils 4.10
Broad regional differences in geology, climate, and vegetation give rise to characteristically different soils. The broadest level of soil classification is the order. Each order has distinctive features. Soil scientists recognize five main soil-forming processes that give rise to these different classes of soils. These processes are laterization, calcification, salinization, podzolization, and gleization.
Soil Erosion Ecological Issues & Applications
Soil erosion on agricultural lands is a serious environmental problem. The removal of natural vegetation and the plowing of the soil destabilizes the soil surface and greatly enhances erosion from wind and water. Sustainable practices such as contour and no-till farming can greatly reduce rates of soil loss.
