Prepare: Measuring leaf area is an important method in plant ecology and agricultural science. Leaf area index (LAI) is one of the most widely measured variables used to describe plant canopies and plant growth rates. In the past, measuring LAI was destructive because every leaf within a given area needed to be collected and measured. Today, there are a variety of simple and advanced methods that can be used to non-destructively estimate LAI.Review: View the following videos and websites to learn about how LAI is measured and how it can be used in ecological, agricultural, and environmental studies.Write: Your post will be based on the alphabetical order of your last names and should be between 250-300 words.Everyone: Explain LAI and provide two examples of how it can be used in environmental research and management. Reference two scholarly or credible sources in addition to your textbook to develop your post. You are encouraged to review the recommended materials this week to help you develop knowledge on the methods and applications of LAI.MY LAST NAME BEGINS WITH A K SO THIS IS WHAT MUST BE DONEIf your last name begins with the letters A through L, then review the video on measuring leaf area with Adobe Photoshop or Image J, and summarize the key points of the method to the class using your own words. Indicate the benefits and limitations of using this method. Use supporting examples to illustrate your points.Measuring Leaf Area with Adobe Photoshop (Links to an external site.)Links to an external site.How to measure Leaf Area (Links to an external site.)Links to an external site.ATTACHED ARE THE WEEKLY RESOURCES TO BE USED PLEASE USE THEM AND DO IT RIGHT THE FIRST TIME SO AS TO AVOID COUNTLESS REVISIONS. MUST USE TWO (2) SCHOLARLY RESOURCES IN ADDITION TO THE TXT FOR A TOTAL OF 3. ATTACHED ARE TO FOUR CHAPTERS FOUR THIS WEEK SO MAKE SURE TO USE THEM ALSO.
How to Measure and Use Leaf Area Index
How to Measure and Use Leaf Area Index
CHAPTER 1 Smith, T. M., & Smith, R. L. (2015). Elements of Ecology (9th ed.). Boston, MA: Pearson. A little more than a year later, on April 22, 1970, as many as 20 million Americans participated in environmental rallies, demonstrations, and other activities as part of the first Earth Day. The New York Times commented on the astonishing rise in environmental awareness, stating that “Rising concern about the environmental crisis is sweeping the nation’s campuses with an intensity that may be on its way to eclipsing student discontent over the war in Vietnam.” Now, more than four decades later, the human population has nearly doubled (3.7 billion in 1970; 7.2 billion as of 2014). Ever-growing demand for basic resources such as food and fuel has created a new array of environmental concerns: resource use and environmental sustainability, the declining biological diversity of our planet, and the potential for human activity to significantly change Earth’s climate. The environmental movement born in the 1970s continues today, and at its core is the belief in the need to redefine our relationship with nature. To do so requires an understanding of nature, and ecology is the particular field of study that provides that understanding. 1.1 Ecology Is the Study of the Relationship between Organisms and Their Environment With the growing environmental movement of the late 1960s and early 1970s, ecology—until then familiar only to a relatively small number of academic and applied biologists—was suddenly thrust into the limelight (see this chapter, Ecological Issues & Applications). Hailed as a framework for understanding the relationship of humans to their environment, ecology became a household word that appeared in newspapers, magazines, and books—although the term was often misused. Even now, people confuse it with terms such as environment and environmentalism. Ecology is neither. Environmentalism is activism with a stated aim of protecting the natural environment, particularly from the negative impacts of human activities. This activism often takes the form of public education programs, advocacy, legislation, and treaties. So what is ecology? Ecology is a science. According to one accepted definition, ecology is the scientific study of the relationships between organisms and their environment. That definition is satisfactory so long as one considers relationships and environment in their fullest meanings. Environment includes the physical and chemical conditions as well as the biological or living components of an organism’s surroundings. Relationships include interactions with the physical world as well as with members of the same and other species. The term ecology comes from the Greek words oikos, meaning “the family household,” and logy, meaning “the study of.” It has the same root word as economics, meaning “management of the household.” In fact, the German zoologist Ernst Haeckel, who originally coined the term ecology in 1866, made explicit reference to this link when he wrote: By ecology we mean the body of knowledge concerning the economy of nature—the investigation of the total relations of the animal both to its inorganic and to its organic; including above all, its friendly and inimical relations with those animals and plants with which it comes directly or indirectly into contact—in a word, ecology is the study of all those complex interrelationships referred to by Darwin as the conditions of the struggle for existence. Haeckel’s emphasis on the relation of ecology to the new and revolutionary ideas put forth in Charles Darwin’s The Origin of Species (1859) is important. Darwin’s theory of natural selection (which Haeckel called “the struggle for existence”) is a cornerstone of the science of ecology. It is a mechanism allowing the study of ecology to go beyond descriptions of natural history and examine the processes that control the distribution and abundance of organisms. 1.2 Organisms Interact with the Environment in the Context of the Ecosystem Organisms interact with their environment at many levels. The physical and chemical conditions surrounding an organism—such as ambient temperature, moisture, concentrations of oxygen and carbon dioxide, and light intensity—all influence basic physiological processes crucial to survival and growth. An organism must acquire essential resources from the surrounding environment, and in doing so, must protect itself from becoming food for other organisms. It must recognize friend from foe, differentiating between potential mates and possible predators. All of this effort is an attempt to succeed at the ultimate goal of all living organisms: to pass their genes on to successive generations. The environment in which each organism carries out this struggle for existence is a place—a physical location in time and space. It can be as large and as stable as an ocean or as small and as transient as a puddle on the soil surface after a spring rain. This environment includes both the physical conditions and the array of organisms that coexist within its confines. This entity is what ecologists refer to as the ecosystem. Organisms interact with the environment in the context of the ecosystem. The eco– part of the word relates to the environment. The –system part implies that the ecosystem functions as a collection of related parts that function as a unit. The automobile engine is an example of a system: components, such as the ignition and fuel pump, function together within the broader context of the engine. Likewise, the ecosystem consists of interacting components that function as a unit. Broadly, the ecosystem consists of two basic interacting components: the living, or biotic, and the nonliving (physical and chemical), or abiotic. Consider a natural ecosystem, such as a forest (Figure 1.2). The physical (abiotic) component of the forest consists of the atmosphere, climate, soil, and water. The biotic component includes the many different organisms—plants, animals, and microbes—that inhabit the forest. Relationships are complex in that each organism not only responds to the abiotic environment but also modifies it and, in doing so, becomes part of the broader environment itself. The trees in the canopy of a forest intercept the sunlight and use this energy to fuel the process of photosynthesis. As a result, the trees modify the environment of the plants below them, reducing the sunlight and lowering air temperature. Birds foraging on insects in the litter layer of fallen leaves reduce insect numbers and modify the environment for other organisms that depend on this shared food resource. By reducing the populations of insects they feed on, the birds are also indirectly influencing the interactions among different insect species that inhabit the forest floor. We will explore these complex interactions between the living and the nonliving environment in greater detail in succeeding chapters. 1.3 Ecological Systems Form a Hierarchy The various kinds of organisms that inhabit our forest make up populations. The term population has many uses and meanings in other fields of study. In ecology, a population is a group of individuals of the same species that occupy a given area. Populations of plants and animals in an ecosystem do not function independently of one another. Some populations compete with other populations for limited resources, such as food, water, or space. In other cases, one population is the food resource for another. Two populations may mutually benefit each other, each doing better in the presence of the other. All populations of different species living and interacting within an ecosystem are referred to collectively as a community. We can now see that the ecosystem, consisting of the biotic community and the abiotic environment, has many levels (Figure 1.3). On one level, individual organisms both respond to and influence the abiotic environment. At the next level, individuals of the same species form populations, such as a population of white oak trees or gray squirrels within a forest. Further, individuals of these populations interact among themselves and with individuals of other species to form a community. Herbivores consume plants, predators eat prey, and individuals compete for limited resources. When individuals die, other organisms consume and break down their remains, recycling the nutrients contained in their dead tissues back into the soil. Organisms interact with the environment in the context of the ecosystem, yet all communities and ecosystems exist in the broader spatial context of the landscape—an area of land (or water) composed of a patchwork of communities and ecosystems. At the spatial scale of the landscape, communities and ecosystems are linked through such processes as the dispersal of organisms and the exchange of materials and energy. Although each ecosystem on the landscape is distinct in that it is composed of a unique combination of physical conditions (such as topography and soils) and associated sets of plant and animal populations (communities), the broad-scale patterns of climate and geology characterizing our planet give rise to regional patterns in the geographic distribution of ecosystems (see Chapter 2). Geographic regions having similar geological and climatic conditions (patterns of temperature, precipitation, and seasonality) support similar types of communities and ecosystems. For example, warm temperatures, high rates of precipitation, and a lack of seasonality characterize the world’s equatorial regions. These warm, wet conditions year-round support vigorous plant growth and highly productive, evergreen forests known as tropical rain forests (see Chapter 23). The broad-scale regions dominated by similar types of ecosystems, such as tropical rain forests, grasslands, and deserts, are referred to as biomes. The highest level of organization of ecological systems is the biosphere—the thin layer surrounding the Earth that supports all of life. In the context of the biosphere, all ecosystems, both on land and in the water, are linked through their interactions—exchanges of materials and energy—with the other components of the Earth system: atmosphere, hydrosphere, and geosphere. Ecology is the study of the complex web of interactions between organisms and their environment at all levels of organization—from the individual organism to the biosphere. 1.4 Ecologists Study Pattern and Process at Many Levels As we shift our focus across the different levels in the hierarchy of ecological systems—from the individual organism to the biosphere—a different and unique set of patterns and processes emerges, and subsequently a different set of questions and approaches for studying these patterns and processes is required (see Figure 1.3). The result is that the broader science of ecology is composed of a range of subdisciplines—from physiological ecology, which focuses on the functioning of individual organisms, to the perspective of Earth’s environment as an integrated system forming the basis of global ecology. Ecologists who focus on the level of the individual examine how features of morphology (structure), physiology, and behavior influence that organism’s ability to survive, grow, and reproduce in its environment. Conversely, how do these same characteristics (morphology, physiology, and behavior) function to constrain the organism’s ability to function successfully in other environments? By contrasting the characteristics of different species that occupy different environments, these ecologists gain insights into the factors influencing the distribution of species. At the individual level, birth and death are discrete events. Yet when we examine the collective of individuals that make up a population, these same processes are continuous as individuals are born and die. At the population level, birth and death are expressed as rates, and the focus of study shifts to examining the numbers of individuals in the population and how these numbers change through time. Populations also have a distribution in space, leading to such questions as how are individuals spatially distributed within an area, and how do the population’s characteristics (numbers and rates of birth and death) change from location to location? As we expand our view of nature to include the variety of plant and animal species that occupy an area, the ecological community, a new set of patterns and processes emerges. At this level of the hierarchy, the primary focus is on factors influencing the relative abundances of various species coexisting within the community. What is the nature of the interactions among the species, and how do these interactions influence the dynamics of the different species’ populations? The diversity of organisms comprising the community modify as well as respond to their surrounding physical environment, and so together the biotic and abiotic components of the environment interact to form an integrated system—the ecosystem. At the ecosystem level, the emphasis shifts from species to the collective properties characterizing the flow of energy and nutrients through the combined physical and biological system. At what rate are energy and nutrients converted into living tissues (termed biomass)? In turn, what processes govern the rate at which energy and nutrients in the form of organic matter (living and dead tissues) are broken down and converted into inorganic forms? What environmental factors limit these processes governing the flow of energy and nutrients through the ecosystem? As we expand our perspective even further, the landscape may be viewed as a patchwork of ecosystems whose boundaries are defined by distinctive changes in the underlying physical environment or species composition. At the landscape level, questions focus on identifying factors that give rise to the spatial extent and arrangement of the various ecosystems that make up the landscape, and ecologists explore the consequences of these spatial patterns on such processes as the dispersal of organisms, the exchange of energy and nutrients between adjacent ecosystems, and the propagation of disturbances such as fire or disease. At a continental to global scale, the questions focus on the broad-scale distribution of different ecosystem types or biomes. How do patterns of biological diversity (the number of different types of species inhabiting the ecosystem) vary geographically across the different biomes? Why do tropical rain forests support a greater diversity of species than do forest ecosystems in the temperate regions? What environmental factors determine the geographic distribution of the different biome types (e.g., forest, grassland, and desert)? Finally, at the biosphere level, the emphasis is on the linkages between ecosystems and other components of the earth system, such as the atmosphere. For example, how does the exchange of energy and materials between terrestrial ecosystems and the atmosphere influence regional and global climate patterns? Certain processes, such as movement of the element carbon between ecosystems and the atmosphere, operate at a global scale and require ecologists to collaborate with oceanographers, geologists, and atmospheric scientists. Throughout our discussion, we have used this hierarchical view of nature and the unique set of patterns and process associated with each level—the individual population, community, ecosystem, landscape, biome, and biosphere—as an organizing framework for studying the science of ecology. In fact, the science of ecology is functionally organized into subdisciplines based on these different levels of organization, each using an array of specialized approaches and methodologies to address the unique set of questions that emerge at these different levels of ecological organization. The patterns and processes at these different levels of organization are linked, however, and identifying these linkages is our objective. For example, at the individual organism level, characteristics such as size, longevity, age at reproduction, and degree of parental care will directly influence rates of birth and survival for the collective of individuals comprising the species’ population. At the community level, the same population will be influenced both positively and negatively through its interactions with populations of other species. In turn, the relative mix of species that make up the community will influence the collective properties of energy and nutrient exchange at the ecosystem level. As we shall see, patterns and processes at each level—from individuals to ecosystems—are intrinsically linked in a web of cause and effect with the patterns and processes operating at the other levels of this organizational hierarchy. 1.5 Ecologists Investigate Nature Using the Scientific Method Although each level in the hierarchy of ecological systems has a unique set of questions on which ecologists focus their research, all ecological studies have one thing in common: they include the process known as the scientific method (Figure 1.4). This method demonstrates the power and limitations of science, and taken individually, each step of the scientific method involves commonplace procedures. Yet taken together, these procedures form a powerful tool for understanding nature. All science begins with observation. In fact, this first step in the process defines the domain of science: if something cannot be observed, it cannot be investigated by science. The observation need not be direct, however. For example, scientists cannot directly observe the nucleus of an atom, yet its structure can be explored indirectly through a variety of methods. Secondly, the observation must be repeatable—able to be made by multiple observers. This constraint helps to minimize unsuspected bias, when an individual might observe what they want or think they ought to observe. The second step in the scientific method is defining a problem—forming a question regarding the observation that has been made. For example, an ecologist working in the prairie grasslands of North America might observe that the growth and productivity (the rate at which plant biomass is being produced per unit area per unit time: grams per meter squared per year [g/m2/yr]) of grasses varies across the landscape. From this observation the ecologist may formulate the question, what environmental factors result in the observed variations in grassland productivity across the landscape? The question typically focuses on seeking an explanation for the observed patterns. Once a question (problem) has been established, the next step is to develop a hypothesis. A hypothesis is an educated guess about what the answer to the question may be. The process of developing a hypothesis is guided by experience and knowledge, and it should be a statement of cause and effect that can be tested. For example, based on her knowledge that nitrogen availability varies across the different soil types found in the region and that nitrogen is an important nutrient limiting plant growth, the ecologist might hypothesize that the observed variations in the growth and productivity of grasses across the prairie landscape are a result of differences in the availability of soil nitrogen. As a statement of cause and effect, certain predictions follow from the hypothesis. If soil nitrogen is the factor limiting the growth and productivity of plants in the prairie grasslands, then grass productivity should be greater in areas with higher levels of soil nitrogen than in areas with lower levels of soil nitrogen. The next step is testing the hypothesis to see if the predictions that follow from the hypothesis do indeed hold true. Interpreting Ecological Data Q1. In the above graph, which variable is the independent variable? Which is the dependent variable? Why? Q2. Would you describe the relationship between available nitrogen and grassland productivity as positive or negative (inverse)? To test this hypothesis, the ecologist may gather data in several ways. The first approach might be a field study to examine how patterns of soil nitrogen and grass productivity covary (vary together) across the landscape. If nitrogen is controlling grassland productivity, productivity should increase with increasing soil nitrogen. The ecologist would measure nitrogen availability and grassland productivity at various sites across the landscape. Then, the relationship between these two variables, nitrogen and productivity, could be expressed graphically (see Quantifying Ecology 1.2 on pages 8 and 9 to learn more about working with graphical data). Visit MasteringBiology at www.masteringbiology.com to work with histograms and scatter plots. After you’ve become familiar with scatter plots, you’ll see the graph of Figure 1.5 shows nitrogen availability on the horizontal or x-axis and grassland productivity on the vertical or y-axis. This arrangement is important. The scientist is assuming that nitrogen is the cause and that grassland productivity is the effect. Because nitrogen (x) is the cause, we refer to it as the independent variable. Because it is hypothesized that grassland productivity (y) is influenced by the availability of nitrogen, we refer to it as the dependent variable. Visit MasteringBiology at www.masteringbiology.com for a tutorial on reading and interpreting graphs. From the observations plotted in Figure 1.5, it is apparent that grassland productivity does, in fact, increase with increasing availability of nitrogen in the soil. Therefore, the data support the hypothesis. Had the data shown no relationship between grassland productivity and nitrogen, the ecologist would have rejected the hypothesis and sought a new explanation for the observed differences in grassland productivity across the landscape. However, although the data suggest that grassland production does increase with increasing soil nitrogen, they do not prove that nitrogen is the only factor controlling grass growth and production. Some other factor that varies with nitrogen availability, such as soil moisture or acidity, may actually be responsible for the observed relationship. To test the hypothesis another way, the ecologist may choose to do an experiment. An experiment is a test under controlled conditions performed to examine the validity of a hypothesis. In designing the experiment, the scientist will try to isolate the presumed causal agent—in this case, nitrogen availability. Quantifying Ecology 1.1 Classifying Ecological Data All ecological studies involve collecting data that includes observations and measurements for testing hypotheses and drawing conclusions about a population. The term population in this context refers to a statistical population. An investigator is highly unlikely to gather observations on all members of a total population, so the part of the population actually observed is referred to as a sample. From this sample data, the investigator will draw her conclusions about the population as a whole. However, not all data are of the same type; and the type of data collected in a study directly influences the mode of presentation, types of analyses that can be performed, and interpretations that can be made. At the broadest level, data can be classified as either categorical or numerical. Categorical data are qualitative, that is, observations that fall into separate and distinct categories. The resulting data are labels or categories, such as the color of hair or feathers, sex, or reproductive status (pre-reproductive, reproductive, post-reproductive). Categorical data can be further subdivided into two categories: nominal and ordinal. Nominal data are categorical data in which objects fall into unordered categories, such as the previous examples of hair color or sex. In contrast, ordinal data are categorical data in which order is important, such as the example of reproductive status. In the special case where only two categories exist, such as in the case of presence or absence of a trait, categorical data are referred to as binary. Both nominal and ordinal data can be binary. With numerical data, objects are “measured” based on some quantitative trait. The resulting data are a set of numbers, such as height, length, or weight. Numerical data can be subdivided into two categories: discrete and continuous. For discrete data, only certain values are possible, such as with integer values or counts. Examples include the number of offspring, number of seeds produced by a plant, or number of times a hummingbird visits a flower during the course of a day. With continuous data, any value within an interval theoretically is possible, limited only by the ability of the measurement device. Examples of this type of data include height, weight, or concentration. What type of data does the variable “available N” (the x-axis) represent in Figure 1.5? How might you transform this variable (available nitrogen) into categorical data? Would it be considered ordinal or nominal? Quantifying Ecology 1.1 Classifying Ecological Data All ecological studies involve collecting data that includes observations and measurements for testing hypotheses and drawing conclusions about a population. The term population in this context refers to a statistical population. An investigator is highly unlikely to gather observations on all members of a total population, so the part of the population actually observed is referred to as a sample. From this sample data, the investigator will draw her conclusions about the population as a whole. However, not all data are of the same type; and the type of data collected in a study directly influences the mode of presentation, types of analyses that can be performed, and interpretations that can be made. At the broadest level, data can be classified as either categorical or numerical. Categorical data are qualitative, that is, observations that fall into separate and distinct categories. The resulting data are labels or categories, such as the color of hair or feathers, sex, or reproductive status (pre-reproductive, reproductive, post-reproductive). Categorical data can be further subdivided into two categories: nominal and ordinal. Nominal data are categorical data in which objects fall into unordered categories, such as the previous examples of hair color or sex. In contrast, ordinal data are categorical data in which order is important, such as the example of reproductive status. In the special case where only two categories exist, such as in the case of presence or absence of a trait, categorical data are referred to as binary. Both nominal and ordinal data can be binary. With numerical data, objects are “measured” based on some quantitative trait. The resulting data are a set of numbers, such as height, length, or weight. Numerical data can be subdivided into two categories: discrete and continuous. For discrete data, only certain values are possible, such as with integer values or counts. Examples include the number of offspring, number of seeds produced by a plant, or number of times a hummingbird visits a flower during the course of a day. With continuous data, any value within an interval theoretically is possible, limited only by the ability of the measurement device. Examples of this type of data include height, weight, or concentration. What type of data does the variable “available N” (the x-axis) represent in Figure 1.5? How might you transform this variable (available nitrogen) into categorical data? Would it be considered ordinal or nominal? The scientist may decide to do a field experiment (Figure 1.6), adding nitrogen to some field sites and not to others. The investigator controls the independent variable (levels of nitrogen) in a predetermined way, to reflect observed variations in soil nitrogen availability across the landscape, and monitors the response of the dependent variable (plant growth). By observing the differences in productivity between the grasslands fertilized with nitrogen and those that were not, the investigator tries to test whether nitrogen is the causal agent. However, in choosing the experimental sites, the ecologist must try to locate areas where other factors that may influence productivity, such as moisture and acidity, are similar. Otherwise, she cannot be sure which factor is responsible for the observed differences in productivity among the sites. Finally, the ecologist might try a third approach—a series of laboratory experiments (Figure 1.7). Laboratory experiments give the investigator much more control over the environmental conditions. For example, she can grow the native grasses in the greenhouse under conditions of controlled temperature, soil acidity, and water availability. If the plants exhibit increased growth with higher nitrogen fertilization, the investigator has further evidence in support of the hypothesis. Nevertheless, she faces a limitation common to all laboratory experiments; that is, the results are not directly applicable in the field. The response of grass plants under controlled laboratory conditions may not be the same as their response under natural conditions in the field. There, the plants are part of the ecosystem and interact with other plants, animals, and the physical environment. Despite this limitation, the ecologist has accumulated additional data describing the basic growth response of the plants to nitrogen availability. Once the observations have been grouped into categories, the resulting frequency distribution can then be displayed as a histogram (type of bar graph; Figure 1a). The x-axis represents the discrete intervals of body length, and the y-axis represents the number of individuals whose body length falls within each given interval. In effect, the continuous data are transformed into categorical data for the purposes of graphical display. Unless there are previous reasons for defining categories, defining intervals is part of the data interpretation process and the search for patterns. For example, how would the pattern represented by the histogram in Figure 1a differ if the intervals were in units of 1 but started with 7.50 (7.50–8.49, 8.50–9.49, etc.)? Often, however, the researcher is examining the relationship between two variables or sets of observations. When both variables are numerical, the most common method of graphically displaying the data is by using a scatter plot. A scatter plot is constructed by defining two axes (x and y), each representing one of the two variables being examined. For example, suppose the researcher who collected the observations of body length for sunfish netted from the pond also measured their weight in grams. The investigator might be interested in whether there is a relationship between body length and weight in sunfish. In this example, body length would be the x-axis, or independent variable (Section 1.5), and body weight would be the y-axis, or dependent variable. Once the two axes are defined, each individual (sunfish) can be plotted as a point on the graph, with the position of the point being defined by its respective values of body length and weight Scatter plots can be described as belonging to one of three general patterns, as shown in Figure 2. In plot (a) there is a general trend for y to increase with increasing values of x. In this case the relationship between x and y is said to be positive (as with the example of body length and weight for sunfish). In plot (b) the pattern is reversed, and y decreases with increasing values of x. In this case the relationship between x and y is said to be negative, or inverse. In plot (c) there is no apparent relationship between x and y. You will find many types of graphs throughout our discussion but most will be histograms and scatter plots. No matter which type of graph is presented, ask yourself the same set of questions—listed below—to help interpret the results. Review this set of questions by applying them to the graphs in Figure 1. What do you find out? What type of data do the observations represent? What variables do each of the axes represent, and what are their units (cm, g, color, etc.)? How do values of y (the dependent variable) vary with values of x (the independent variable)? Go to Analyzing Ecological Data at www.masteringbiology.com to further explore how to display data graphically. Having conducted several experiments that confirm the link between patterns of grass productivity to nitrogen availability, the ecologist may now wish to explore this relationship further, to see how the relationship between productivity and nitrogen is influenced by other environmental factors that vary across the prairie landscape. For example, how do differences in rainfall and soil moisture across the region influence the relationship between grass production and soil nitrogen? Once again hypotheses are developed, predictions made, and experiments conducted. As the ecologist develops a more detailed understanding of how various environmental factors interact with soil nitrogen to control grass production, a more general theory of the influence of environmental factors controlling grass production in the grassland prairies may emerge. A theory is an integrated set of hypotheses that together explain a broader set of observations than any single hypothesis—such as a general theory of environmental controls on productivity of the prairie grassland ecosystems of North America. Although the diagram of the scientific method presented in Figure 1.4 represents the process of scientific investigation as a sequence of well-defined steps that proceeds in a linear fashion, in reality, the process of scientific research often proceeds in a nonlinear fashion. Scientists often begin an investigation based on readings of previously published studies, discussions with colleagues, or informal observations made in the field or laboratory rather than any formal process. Often during hypothesis testing, observations may lead the researcher to modify the experimental design or redefine the original hypothesis. In reality, the practice of science involves unexpected twist and turns. In some cases, unexpected observations or results during the initial investigation may completely change the scope of the study, leading the researcher in directions never anticipated. Whatever twists and unanticipated turns may occur, however, the process of science is defined by the fundamental structure and constraints of the scientific method. 1.6 Models Provide a Basis for Predictions Scientists use the understanding derived from observation and experiments to develop models. Data are limited to the special case of what happened when the measurements were made. Like photographs, data represent a given place and time. Models use the understanding gained from the data to predict what will happen in some other place and time. Models are abstract, simplified representations of real systems. They allow us to predict some behavior or response using a set of explicit assumptions, and as with hypotheses, these predictions should be testable through further observation or experiments. Models may be mathematical, like computer simulations, or they may be verbally descriptive, like Darwin’s theory of evolution by natural selection (see Chapter 5). Hypotheses are models, although the term model is typically reserved for circumstances in which the hypothesis has at least some limited support through observations and experimental results. For example, the hypothesis relating grass production to nitrogen availability is a model. It predicts that plant productivity will increase with increasing nitrogen availability. However, this prediction is qualitative—it does not predict how much plant productivity will increase. In contrast, mathematical models usually offer quantitative predictions. For example, from the data in Figure 1.5, we can develop a regression equation—a form of statistical model—to predict the amount of grassland productivity per unit of nitrogen in the soil Interpreting Ecological Data Q1. How could you use the simple linear regression model presented to predict productivity for a grassland site not included in the graph? Q2. What is the predicted productivity for a site with available nitrogen of 5 g/m2/yr? (Use the linear regression equation.) All of the approaches just discussed—observation, experimentation, hypothesis testing, and development of models—appear throughout our discussion to illustrate basic concepts and relationships. They are the basic tools of science. For every topic, an array of figures and tables present the observations, experimental data, and model predictions used to test specific hypotheses regarding pattern and process at the different levels of ecological organization. Being able to analyze and interpret the data presented in these figures and tables is essential to your understanding of the science of ecology. To help you develop these skills, we have annotated certain figures and tables to guide you in their interpretation. In other cases, we pose questions that ask you to interpret, analyze, and draw conclusions from the data presented. These figures and tables are labeled Interpreting Ecological Data. (See Figure 2.16 on page 23 for the first example.) 1.7 Uncertainty Is an Inherent Feature of Science Collecting observations, developing and testing hypotheses, and constructing predictive models all form the backbone of the scientific method (see Figure 1.4). It is a continuous process of testing and correcting concepts to arrive at explanations for the variation we observe in the world around us, thus unifying observations that on first inspection seem unconnected. The difference between science and art is that, although both pursuits involve creation of concepts, in science, the exploration of concepts is limited to the facts. In science, the only valid means of judging a concept is by testing its empirical truth. However, scientific concepts have no permanence because they are only our interpretations of natural phenomena. We are limited to inspecting only a part of nature because to understand, we have to simplify. As discussed in Section 1.5, in designing experiments, we control the pertinent factors and try to eliminate others that may confuse the results. Our intent is to focus on a subset of nature from which we can establish cause and effect. The trade-off is that whatever cause and effect we succeed in identifying represents only a partial connection to the nature we hope to understand. For that reason, when experiments and observations support our hypotheses, and when the predictions of the models are verified, our job is still not complete. We work to loosen the constraints imposed by the need to simplify so that we can understand. We expand our hypothesis to cover a broader range of conditions and once again begin testing its ability to explain our new observations. It may sound odd at first, but science is a search for evidence that proves our concepts wrong. Rarely is there only one possible explanation for an observation. As a result, any number of hypotheses may be developed that might be consistent with an observation. The determination that experimental data are consistent with a hypothesis does not prove that the hypothesis is true. The real goal of hypothesis testing is to eliminate incorrect ideas. Thus, we must follow a process of elimination, searching for evidence that proves a hypothesis wrong. Science is essentially a self-correcting activity, dependent on the continuous process of debate. Dissent is the activity of science, fueled by free inquiry and independence of thought. To the outside observer, this essential process of debate may appear to be a shortcoming. After all, we depend on science for the development of technology and the ability to solve problems. For the world’s current environmental issues, the solutions may well involve difficult ethical, social, and economic decisions. In this case, the uncertainty inherent in science is discomforting. However, we must not mistake uncertainty for confusion, nor should we allow disagreement among scientists to become an excuse for inaction. Instead, we need to understand the uncertainty so that we may balance it against the costs of inaction. 1.8 Ecology Has Strong Ties to Other Disciplines The complex interactions taking place within ecological systems involve all kinds of physical, chemical, and biological processes. To study these interactions, ecologists must draw on other sciences. This dependence makes ecology an interdisciplinary science. Although we explore topics that are typically the subject of disciplines such as biochemistry, physiology, and genetics, we do so only in the context of understanding the interplay of organisms with their environment. The study of how plants take up carbon dioxide and lose water, for example, belongs to plant physiology (see Chapter 6). Ecology looks at how these processes respond to variations in rainfall and temperature. This information is crucial to understanding the distribution and abundance of plant populations and the structure and function of ecosystems on land. Likewise, we must draw on many of the physical sciences, such as geology, hydrology, and meteorology. They help us chart other ways in which organisms and environments interact. For instance, as plants take up water, they influence soil moisture and the patterns of surface water flow. As they lose water to the atmosphere, they increase atmospheric water content and influence regional patterns of precipitation. The geology of an area influences the availability of nutrients and water for plant growth. In each example, other scientific disciplines are crucial to understanding how individual organisms both respond to and shape their environment. In the 21st century, ecology is entering a new frontier, one that requires expanding our view of ecology to include the dominant role of humans in nature. Among the many environmental problems facing humanity, four broad and interrelated areas are crucial: human population growth, biological diversity, sustainability, and global climate change. As the human population increased from approximately 500 million to more than 7 billion in the past two centuries, dramatic changes in land use have altered Earth’s surface. The clearing of forests for agriculture has destroyed many natural habitats, resulting in a rate of species extinction that is unprecedented in Earth’s history. In addition, the expanding human population is exploiting natural resources at unsustainable levels. As a result of the growing demand for energy from fossil fuels that is needed to sustain economic growth, the chemistry of the atmosphere is changing in ways that are altering Earth’s climate. These environmental problems are ecological in nature, and the science of ecology is essential to understanding their causes and identifying ways to mitigate their impacts. Addressing these issues, however, requires a broader interdisciplinary framework to better understand their historical, social, legal, political, and ethical dimensions. That broader framework is known as environmental science. Environmental science examines the impact of humans on the natural environment and as such covers a wide range of topics including agronomy, soils, demography, agriculture, energy, and hydrology, to name but a few. Throughout the text, we use the Ecological Issues & Applications sections of each chapter to highlight topics relating to current environmental issues regarding human impacts on the environment and to illustrate the importance of the science of ecology to better understanding the human relationship with the environment. 1.9 The Individual Is the Basic Unit of Ecology As we noted previously, ecology encompasses a broad area of investigation—from the individual organism to the biosphere. Our study of the science of ecology uses this hierarchical framework in the chapters that follow. We begin with the individual organism, examining the processes it uses and constraints it faces in maintaining life under varying environmental conditions. The individual organism forms the basic unit in ecology. The individual senses and responds to the prevailing physical environment. The collective properties of individual births and deaths drive the dynamics of populations, and individuals of different species interact with one another in the context of the community. But perhaps most importantly, the individual, through the process of reproduction, passes genetic information to successive individuals, defining the nature of individuals that will compose future populations, communities, and ecosystems. At the individual level we can begin to understand the mechanisms that give rise to the diversity of life and ecosystems on Earth—mechanisms that are governed by the process of natural selection. But before embarking on our study of ecological systems, we examine characteristics of the abiotic (physical and chemical) environment that function to sustain and constrain the patterns of life on our planet. Ecological Issues & Applications Ecology Has a Rich History The genealogy of most sciences is direct. Tracing the roots of chemistry and physics is relatively easy. The science of ecology is different. Its roots are complex and intertwined with a wide array of scientific advances that have occurred in other disciplines within the biological and physical sciences. Although the term ecology did not appear until the mid-19th century and took another century to enter the vernacular, the idea of ecology is much older. Arguably, ecology goes back to the ancient Greek scholar Theophrastus, a friend of Aristotle, who wrote about the relations between organisms and the environment. On the other hand, ecology as we know it today has vital roots in plant geography and natural history. In the 1800s, botanists began exploring and mapping the world’s vegetation. One of the early plant geographers was Carl Ludwig Willdenow (1765–1812). He pointed out that similar climates supported vegetation similar in form, even though the species were different. Another was Friedrich Heinrich Alexander von Humboldt (1769–1859), for whom the Humboldt Current, flowing along the west coast of South America, is named. He spent five years exploring Latin America, including the Orinoco and Amazon rivers. Humboldt correlated vegetation with environmental characteristics and coined the term plant association. The recognition that the form and function of plants within a region reflects the constraints imposed by the physical environment led the way for a new generation of scientists that explored the relationship between plant biology and plant geography (see Chapter 23). Among this new generation of plant geographers was Johannes Warming (1841–1924) at the University of Copenhagen, who studied the tropical vegetation of Brazil. He wrote the first text on plant ecology, Plantesamfund. Warming integrated plant morphology, physiology, taxonomy, and biogeography into a coherent whole. This book had a tremendous influence on the development of ecology. Meanwhile, activities in other areas of natural history also assumed important roles. One was the voyage of Charles Darwin (1809–1882) on the Beagle. Working for years on notes and collections from this trip, Darwin compared similarities and dissimilarities among organisms within and among continents. He attributed differences to geological barriers. He noted how successive groups of plants and animals, distinct yet obviously related, replaced one another. Developing his theory of evolution and the origin of species, Darwin came across the writings of Thomas Malthus (1766–1834). An economist, Malthus advanced the principle that populations grow in a geometric fashion, doubling at regular intervals until they outstrip the food supply. Ultimately, a “strong, constantly operating force such as sickness and premature death” would restrain the population. From this concept Darwin developed the idea of “natural selection” as the mechanism guiding the evolution of species (see Chapter 5). Meanwhile, unbeknownst to Darwin, an Austrian monk, Gregor Mendel (1822–1884), was studying the transmission of characteristics from one generation of pea plants to another in his garden. Mendel’s work on inheritance and Darwin’s work on natural selection provided the foundation for the study of evolution and adaptation, the field of population genetics. Darwin’s theory of natural selection, combined with the new understanding of genetics (the means by which characteristics are transmitted from one generation to the next) provided the mechanism for understanding the link between organisms and their environment, which is the focus of ecology. Early ecologists, particularly plant ecologists, were concerned with observing the patterns of organisms in nature, and attempting to understand how patterns were formed and maintained by interactions with the physical environment. Some, notably Frederic E. Clements (Figure 1.9), sought some system of organizing nature. He proposed that the plant community behaves as a complex organism or superorganism that grows and develops through stages to a mature or climax state (see Chapter 16). His idea was accepted and advanced by many ecologists. A few ecologists, however, notably Arthur G. Tansley, did not share this view. In its place Tansley advanced a holistic and integrated ecological concept that combined living organisms and their physical environment into a system, which he called the ecosystem (see Chapter 20). Whereas the early plant ecologists were concerned mostly with terrestrial vegetation, another group of European biologists was interested in the relationship between aquatic plants and animals and their environment. They advanced the ideas of organic nutrient cycling and feeding levels, using the terms producers and consumers. Their work influenced a young limnologist at the University of Minnesota, R. A. Lindeman. He traced “energy-available” relationships within a lake community. His 1942 paper, “The Trophic-Dynamic Aspects of Ecology,” marked the beginning of ecosystem ecology, the study of whole living systems. Lindeman’s theory stimulated further pioneering work in the area of energy flow and nutrient cycling by G. E. Hutchinson of Yale University (Figure 1.10) and E. P. and H. T. Odum of the University of Georgia. Their work became a foundation of ecosystem ecology. The use of radioactive tracers, a product of the atomic age, to measure the movements of energy and nutrients through ecosystems and the use of computers to analyze large amounts of data stimulated the development of systems ecology, the application of general system theory and methods to ecology. Animal ecology initially developed largely independently of the early developments in plant ecology. The beginnings of animal ecology can be traced to two Europeans, R. Hesse of Germany and Charles Elton of England. Elton’s Animal Ecology (1927) and Hesse’s Tiergeographie auf logischer grundlage (1924), translated into English as Ecological Animal Geography, strongly influenced the development of animal ecology in the United States. Charles Adams and Victor Shelford were two pioneering U.S. animal ecologists. Adams published the first textbook on animal ecology, A Guide to the Study of Animal Ecology (1913). Shelford wrote Animal Communities in Temperate America (1913). Shelford gave a new direction to ecology by stressing the interrelationship between plants and animals. Ecology became a science of communities. Some previous European ecologists, particularly the marine biologist Karl Mobius, had developed the general concept of the community. In his essay “An Oyster Bank is a Biocenose” (1877), Mobius explained that the oyster bank, although dominated by one animal, was really a complex community of many interdependent organisms. He proposed the word biocenose for such a community. The word comes from the Greek, meaning life having something in common. The appearance in 1949 of the encyclopedic Principles of Animal Ecology by five second-generation ecologists from the University of Chicago (W. C. Allee, A. E. Emerson, Thomas Park, Orlando Park, and K. P. Schmidt) pointed to the direction that modern ecology would take. It emphasized feeding relationships and energy budgets, population dynamics, and natural selection and evolution. During the period of development of the field of animal ecology, natural history observations also focused on the behavior of animals. This focus on animal behavior began with 19th-century behavioral studies including those of ants by William Wheeler and of South American monkeys by Charles Carpenter. Later, the pioneering studies of Konrad Lorenz and Niko Tinbergen on the role of imprinting and instinct in the social life of animals, particularly birds and fish, gave rise to ethology. It spawned an offshoot, behavioral ecology, exemplified by L. E. Howard’s early study on territoriality in birds. Behavioral ecology is concerned with intraspecific and interspecific relationships such as mating, foraging, defense, and how behavior is influenced by natural selection. The writings of the economist Malthus that were so influential in the development of Darwin’s ideas regarding the origin of species also stimulated the study of natural populations. The study of populations in the early 20th century branched into two fields. One, population ecology, is concerned with population growth (including birthrates and death rates), regulation and intraspecific and interspecific competition, mutualism, and predation. The other, a combination of population genetics and population ecology is evolutionary ecology, which deals with the role of natural selection in physical and behavioral adaptations and speciation. Focusing on adaptations, physiological ecology is concerned with the responses of individual organisms to temperature, moisture, light, and other environmental conditions. Closely associated with population and evolutionary ecology is community ecology, with its focus on species interactions. One of the major objectives of community ecology is to understand the origin, maintenance, and consequences of species diversity within ecological communities. With advances in biology, physics, and chemistry throughout the latter part of the 20th century, new areas of study in ecology emerged. The development of aerial photography and later the launching of satellites by the U.S. space program provided scientists with a new perspective of the surface of Earth through the use of remote sensing data. Ecologists began to explore spatial processes that linked adjacent communities and ecosystems through the new emerging field of landscape ecology. A new appreciation of the impact of changing land use on natural ecosystems led to the development of conservation ecology, which applies principles from different fields, from ecology to economics and sociology, to the maintenance of biological diversity. The application of principles of ecosystem development and function to the restoration and management of disturbed lands gave rise to restoration ecology, whereas understanding Earth as a system is the focus of the newest area of ecological study, global ecology. Ecology has so many roots that it probably will always remain multifaceted—as the ecological historian Robert McIntosh calls it, “a polymorphic discipline.” Insights from these many specialized areas of ecology will continue to enrich the science as it moves forward in the 21st century. Summary Ecology 1.1 Ecology is the scientific study of the relationships between organisms and their environment. The environment includes the physical and chemical conditions and biological or living components of an organism’s surroundings. Relationships include interactions with the physical world as well as with members of the same and other species. Ecosystems 1.2 Organisms interact with their environment in the context of the ecosystem. Broadly, the ecosystem consists of two components, the living (biotic) and the physical (abiotic), interacting as a system. Hierarchical Structure 1.3 Ecological systems may be viewed in a hierarchical framework, from individual organisms to the biosphere. Organisms of the same species that inhabit a given physical environment make up a population. Populations of different kinds of organisms interact with members of their own species as well as with individuals of other species. These interactions range from competition for shared resources to interactions that are mutually beneficial for the individuals of both species involved. Interacting populations make up a biotic community. The community plus the physical environment make up an ecosystem. All communities and ecosystems exist in the broader spatial context of the landscape—an area of land (or water) composed of a patchwork of communities and ecosystems. Geographic regions having similar geological and climatic conditions support similar types of communities and ecosystems, referred to as biomes. The highest level of organization of ecological systems is the biosphere—the thin layer around Earth that supports all of life. Ecological Studies 1.4 At each level in the hierarchy of ecological systems—from the individual organism to the biosphere—a different and unique set of patterns and processes emerges; subsequently, a different set of questions and approaches for studying these patterns and processes is required. Scientific Method 1.5 All ecological studies are conducted by using the scientific method. All science begins with observation, from which questions emerge. The next step is the development of a hypothesis—a proposed answer to the question. The hypothesis must be testable through observation and experiments. Models 1.6 From research data, ecologists develop models. Models allow us to predict some behavior or response using a set of explicit assumptions. They are abstractions and simplifications of natural phenomena. Such simplification is necessary to understand natural processes. Uncertainty in Science 1.7 An inherent feature of scientific study is uncertainty; it arises from the limitation posed by focusing on only a small subset of nature, and it results in an incomplete perspective. Because we can develop any number of hypotheses that may be consistent with an observation, determining that experimental data are consistent with a hypothesis is not sufficient to prove that the hypothesis is true. The real goal of hypothesis testing is to eliminate incorrect ideas. An Interdisciplinary Science 1.8 Ecology is an interdisciplinary science because the interactions of organisms with their environment and with one another involve physiological, behavioral, and physical responses. The study of these responses draws on such fields as physiology, biochemistry, genetics, geology, hydrology, and meteorology. Individuals 1.9 The individual organism forms the basic unit in ecology. It is the individual that responds to the environment and passes genes to successive generations. It is the collective birth and death of individuals that determines the dynamics of populations, and the interactions among individuals of the same and different species that structures communities. History Ecological Issues & Applications Ecology has its origin in natural history and plant geography. Over the past century it has developed into a science that has its roots in disciplines as diverse as genetics and systems engineering.
How to Measure and Use Leaf Area Index
CHAPTER 2 Smith, T. M., & Smith, R. L. (2015). Elements of Ecology (9th ed.). Boston, MA: Pearson. 2.1 Surface Temperatures Reflect the Difference between Incoming and Outgoing Radiation Solar radiation—the electromagnetic energy (Figure 2.1) emanating from the Sun—travels more or less unimpeded through the vacuum of space until it reaches Earth’s atmosphere. Scientists conceptualize solar radiation as a stream of photons, or packets of energy, that—in one of the great paradoxes of science—behave either as waves or as particles, depending on how they are observed. Scientists characterize waves of energy in terms of their wavelength (λ), or the physical distances between successive crests, and their frequency (ν), or the number of crests that pass a given point per second. All objects emit radiant energy, typically across a wide range of wavelengths. The exact nature of the energy emitted, however, depends on the object’s temperature (Figure 2.2). The hotter the object is, the more energetic the emitted photons and the shorter the wavelength. A hot surface such as that of the Sun (~5800°C) gives off primarily shortwave (solar) radiation. In contrast, cooler objects such as Earth’s surface (average temperature of 15°C) emit radiation of longer wavelengths, or longwave (terrestrial) radiation. Some of the shortwave radiation that reaches the surface of our planet is reflected back into space. The quantity of shortwave radiation reflected by a surface is a function of its reflectivity, referred to as its albedo. Albedo is expressed as a proportion (0–1.0) of the shortwave radiation striking a surface that is reflected and differs for different surfaces. For example, surfaces covered by ice and snow have a high albedo (0.8–0.9), reflecting anywhere from 80 to 90 percent of incoming solar radiation, whereas a forest has a relatively low albedo (0.05), reflecting only 5 percent of sunlight. The global annual averaged albedo is approximately 0.30 (30 percent reflectance). The difference between the incoming shortwave radiation and the reflected shortwave radiation is the net shortwave radiation absorbed by the surface. In turn, some of the energy absorbed by Earth’s surface (both land and water) is emitted back out into space as terrestrial longwave radiation. The amount of energy emitted is dependent on the temperature of the surface. The hotter the surface, the more radiant energy it will emit. Most of the longwave radiation emitted by Earth’s surface, however, is absorbed by water vapor and carbon dioxide in the atmosphere. This absorbed radiation is emitted downward toward the surface as longwave atmospheric radiation, which keeps near surface temperatures warmer than they would be without this blanket of gases. This is known as the “greenhouse effect,” and gases such as water vapor and carbon dioxide that are good absorbers of longwave radiation are known as “greenhouse gases.” It is the difference between the incoming shortwave (solar) radiation and outgoing longwave (terrestrial) radiation that defines the net radiation (Figure 2.3) and determines surface temperatures. If the amount of incoming shortwave radiation exceeds the amount of outgoing longwave radiation, surface temperature increases. Conversely, surface temperature declines if the quantity of outgoing longwave radiation exceeds the incoming shortwave radiation (as is the case during the night). On average, the amount of incoming shortwave radiation intercepted by Earth and the quantity of longwave radiation emitted by the planet back into space balance, and the average surface temperature of our planet remains approximately 15oC. Note, however, from the global map of average annual surface net radiation presented in (Figure 2.4) that there is a distinct latitudinal gradient of decreasing net surface radiation from the equator toward the poles. This decline is a direct function of the variation with latitude in the amount of shortwave radiation reaching the surface. Two factors influence this variation (Figure 2.5). First, at higher latitudes, solar radiation hits the surface at a steeper angle, spreading sunlight over a larger area. Second, solar radiation that penetrates the atmosphere at a steep angle must travel through a deeper layer of air. In the process, it encounters more particles in the atmosphere, which reflect more of the shortwave radiation back into space. The result of the decline in net radiation with latitude is a distinct gradient of decreasing mean annual temperature from the equator toward the poles (Figure 2.6). Some of the shortwave radiation that reaches the surface of our planet is reflected back into space. The quantity of shortwave radiation reflected by a surface is a function of its reflectivity, referred to as its albedo. Albedo is expressed as a proportion (0–1.0) of the shortwave radiation striking a surface that is reflected and differs for different surfaces. For example, surfaces covered by ice and snow have a high albedo (0.8–0.9), reflecting anywhere from 80 to 90 percent of incoming solar radiation, whereas a forest has a relatively low albedo (0.05), reflecting only 5 percent of sunlight. The global annual averaged albedo is approximately 0.30 (30 percent reflectance). The difference between the incoming shortwave radiation and the reflected shortwave radiation is the net shortwave radiation absorbed by the surface. In turn, some of the energy absorbed by Earth’s surface (both land and water) is emitted back out into space as terrestrial longwave radiation. The amount of energy emitted is dependent on the temperature of the surface. The hotter the surface, the more radiant energy it will emit. Most of the longwave radiation emitted by Earth’s surface, however, is absorbed by water vapor and carbon dioxide in the atmosphere. This absorbed radiation is emitted downward toward the surface as longwave atmospheric radiation, which keeps near surface temperatures warmer than they would be without this blanket of gases. This is known as the “greenhouse effect,” and gases such as water vapor and carbon dioxide that are good absorbers of longwave radiation are known as “greenhouse gases.” It is the difference between the incoming shortwave (solar) radiation and outgoing longwave (terrestrial) radiation that defines the net radiation (Figure 2.3) and determines surface temperatures. If the amount of incoming shortwave radiation exceeds the amount of outgoing longwave radiation, surface temperature increases. Conversely, surface temperature declines if the quantity of outgoing longwave radiation exceeds the incoming shortwave radiation (as is the case during the night). On average, the amount of incoming shortwave radiation intercepted by Earth and the quantity of longwave radiation emitted by the planet back into space balance, and the average surface temperature of our planet remains approximately 15oC. Note, however, from the global map of average annual surface net radiation presented in (Figure 2.4) that there is a distinct latitudinal gradient of decreasing net surface radiation from the equator toward the poles. This decline is a direct function of the variation with latitude in the amount of shortwave radiation reaching the surface. Two factors influence this variation (Figure 2.5). First, at higher latitudes, solar radiation hits the surface at a steeper angle, spreading sunlight over a larger area. Second, solar radiation that penetrates the atmosphere at a steep angle must travel through a deeper layer of air. In the process, it encounters more particles in the atmosphere, which reflect more of the shortwave radiation back into space. The result of the decline in net radiation with latitude is a distinct gradient of decreasing mean annual temperature from the equator toward the poles (Figure 2.6). 2.2 Intercepted Solar Radiation and Surface Temperatures Vary Seasonally Although the variation in shortwave (solar) radiation reaching Earth’s surface with latitude can explain the gradient of decreasing mean annual temperature from the equator to the poles, it does not explain the systematic variation occurring over the course of a year. What gives rise to the seasons on Earth? Why do the hot days of summer give way to the changing colors of fall, or the freezing temperatures and snow-covered landscape of winter to the blanket of green signaling the onset of spring? The explanation is quite simple: it is because Earth does not stand up straight but rather tilts to its side. Earth, like all planets, is subject to two distinct motions. While it orbits the Sun, Earth rotates about an axis that passes through the North and South Poles, giving rise to the brightness of day followed by the darkness of night (the diurnal cycle). Earth travels about the Sun in an ecliptic plane. By chance, Earth’s axis of spin is not perpendicular to the ecliptic plane but tilted at an angle of 23.5°. As a result, as Earth follows its elliptical orbit about the Sun, the location on the surface where the Sun is directly overhead at midday migrates between 23.5° N and 23.5° S latitude over the course of the year (Figure 2.7). At the vernal equinox (approximately March 21) and autumnal equinox (approximately September 22), the Sun is directly overhead at the equator (see Figure 2.7). At this time, the equatorial region receives the greatest input of shortwave (solar) radiation, and every place on Earth receives the same 12 hours each of daylight and night. At the summer solstice (approximately June 22) in the Northern Hemisphere, solar rays fall directly on the Tropic of Cancer (23.5° N; see Figure 2.9). This is when days are longest in the Northern Hemisphere, and the input of solar radiation to the surface is the greatest. In contrast, the Southern Hemisphere experiences winter at this time. At winter solstice (about December 22) in the Northern Hemisphere, solar rays fall directly on the Tropic of Capricorn (23.5° S; see Figure 2.7). This period is summer in the Southern Hemisphere, whereas the Northern Hemisphere is enduring shorter days and colder temperatures. Thus, the summer solstice in the Northern Hemisphere is the winter solstice in the Southern Hemisphere. In the equatorial region there is little seasonality (variation over the year) in net radiation, temperature, or day length. Seasonality systematically increases from the equator to the poles (Figure 2.8). At the Arctic and Antarctic circles (66.5° N and S, respectively), day length varies from 0 to 24 hours over the course of the year. The days shorten until the winter solstice, a day of continuous darkness. The days lengthen with spring, and on the day of the summer solstice, the Sun never sets. 2.3 Geographic Difference in Surface Net Radiation Result in Global Patterns of Atmospheric Circulation As we discussed in the previous section, the average net radiation of the planet is zero; that is to say that the amount of incoming shortwave radiation absorbed by the surface is offset by the quantity of outgoing longwave radiation back into space. Otherwise, the average temperature of the planet would either increase or decrease. Geographically, however, this is not the case. Note from the global map of mean annual net radiation presented in Figure 2.4 that there are regions of positive (surplus) and negative (deficit) net radiation. In fact, there is a distinct latitudinal pattern of surface radiation illustrated in Figure 2.9. Between 35.5° N and 35.5° S (from the equator to the midlatitudes), the amount of incoming shortwave radiation received over the year exceeds the amount of outgoing longwave radiation and there is a surplus. In contrast, from 35.5° N and S latitude to the poles (90° N and S), the amount of outgoing longwave radiation over the year exceeds the incoming shortwave radiation and there is a deficit. This imbalance in net radiation sets into motion a global scale pattern of the redistribution of thermal energy (heat) from the equator to the poles. Recall from basic physical sciences that energy flows from regions of higher concentration to regions of lower concentration, that is, from warmer regions to cooler regions. The primary mechanism of this planetary transfer of heat from the tropics (region of net radiation surplus) to the poles (region of net radiation deficit) is the process of convection, that is, the transfer of heat through the circulation of fluids (air and water). As previously discussed, the equatorial region receives the largest annual input of solar radiation and greatest net radiation surplus. Air warmed at the surface rises because it is less dense than the cooler air above it. Air heated at the equatorial region rises to the top of the troposphere, establishing a zone of low pressure at the surface (Figure 2.10). This low atmospheric pressure at the surface causes air from the north and south to flow toward the equator (air moves from areas of higher pressure to areas of lower pressure). The resulting convergence of winds from the north and south in the region of the equator is called the Intertropical Convergence Zone, or ITCZ, for short. The continuous column of rising air at the equator forces the air mass above to spread north and south toward the poles. As air masses move poleward, they cool, become heavier (more dense), and sink. The sinking air at the poles raises surface air pressure, forming a high-pressure zone and creating a pressure gradient from the poles to the equator. The cooled, heavier air then flows toward the low-pressure zone at the equator, replacing the warm air rising over the tropics and closing the pattern of air circulation. If Earth were stationary and without irregular landmasses, the atmosphere would circulate as shown in Figure 2.10. Earth, however, spins on its axis from west to east. Although each point on Earth’s surface makes a complete rotation every 24 hours, the speed of rotation varies with latitude (and circumference). At a point on the equator (its widest circumference at 40,176 km), the speed of rotation is 1674 km per hour. In contrast, at 60° N or S, Earth’s circumference is approximately half that at the equator (20,130 km), and the speed of rotation is 839 km per hour. According to the law of angular motion, the momentum of an object moving from a greater circumference to a lesser circumference will deflect in the direction of the spin, and an object moving from a lesser circumference to a greater circumference will deflect in the direction opposite that of the spin. As a result, air masses and all moving objects in the Northern Hemisphere are deflected to the right (clockwise motion), and in the Southern Hemisphere to the left (counterclockwise motion). This deflection in the pattern of air flow is the Coriolis effect, named after the 19th-century French mathematician G. C. Coriolis, who first analyzed the phenomenon (Figure 2.11). In addition to the deflection resulting from the Coriolis effect, air that moves poleward is subject to longitudinal compression, that is, poleward-moving air is forced into a smaller space, and the density of the air increases. These factors prevent a direct, simple flow of air from the equator to the poles. Instead, they create a series of belts of prevailing winds, named for the direction they come from. These belts break the simple flow of surface air toward the equator and they flow aloft to the poles into a series of six cells, three in each hemisphere. They produce areas of low and high pressure as air masses ascend from and descend toward the surface, respectively (Figure 2.12). To trace the flow of air as it circulates between the equator and poles, we begin at Earth’s equatorial region, which receives the largest annual input of solar radiation. Air heated in the equatorial zone rises upward, creating a low-pressure zone near the surface—the equatorial low. This upward flow of air is balanced by a flow of air from the north and south toward the equator (ITCZ). As the warm air mass rises, it begins to spread, diverging northward and southward toward the North and South Poles, cooling as it goes. In the Northern Hemisphere, the Coriolis effect forces air in an easterly direction, slowing its progress north. At about 30° N, the now-cool air sinks, closing the first of the three cells—the Hadley cells, named for the Englishman George Hadley, who first described this pattern of circulation in 1735. The descending air forms a semipermanent high-pressure belt at the surface that encircles Earth—the subtropical high. Having descended, the cool air warms and splits into two currents flowing over the surface. One moves northward toward the pole, diverted to the right by the Coriolis effect to become the prevailing westerlies. Meanwhile, the other current moves southward toward the equator. Also deflected to the right, this southward-flowing stream becomes the strong, reliable winds that were called trade winds by the 17th-century merchant sailors who used them to reach the Americas from Europe. In the Northern Hemisphere, these winds are known as the northeast trades. In the Southern Hemisphere, where similar flows take place, these winds are known as the southeast trades. As the mild air of the westerlies moves poleward, it encounters cold air moving down from the pole (approximately 60° N). These two air masses of contrasting temperature do not readily mix. They are separated by a boundary called the polar front—a zone of low pressure (the subpolar low) where surface air converges and rises. Some of the rising air moves southward until it reaches approximately 30° latitude (the region of the subtropical high), where it sinks back to the surface and closes the second of the three cells—the Ferrel cell, named after U.S. meteorologist William Ferrel. As the northward-moving air reaches the pole, it slowly sinks to the surface and flows back (southward) toward the polar front, completing the last of the three cells—the polar cell. This southward-moving air is deflected to the right by the Coriolis effect, giving rise to the polar easterlies. Similar flows occur in the Southern Hemisphere (see Figure 2.12). This pattern of global atmospheric circulation functions to transport heat (thermal energy) from the tropics (the region of net radiation surplus) toward the poles (the regions of net radiation deficit), moderating temperatures at the higher latitudes. 2.4 Surface Winds and Earth’s Rotation Create Ocean Currents The global pattern of prevailing winds plays a crucial role in determining major patterns of surface water flow in Earth’s oceans. These systematic patterns of water movement are called currents. In fact, until they encounter one of the continents, the major ocean currents generally mimic the movement of the surface winds presented in the previous section. Each ocean is dominated by two great circular water motions, or gyres. Within each gyre, the ocean current moves clockwise in the Northern Hemisphere and counterclockwise in the Southern Hemisphere (Figure 2.13). Along the equator, trade winds push warm surface waters westward. When these waters encounter the eastern margins of continents, they split into north- and south-flowing currents along the coasts, forming north and south gyres. As the currents move farther from the equator, the water cools. Eventually, they encounter the westerly winds at higher latitudes (30–60° N and 30–60° S), which produce eastward-moving currents. When these eastward-moving currents encounter the western margins of the continents, they form cool currents that flow along the coastline toward the equator. Just north of the Antarctic continent, ocean waters circulate unimpeded around the globe. As with the patterns of global atmospheric circulation and winds, the gyres function to redistribute heat from the tropics northward and southward toward the poles 2.5 Temperature Influences the Moisture Content of Air Air temperature plays a crucial role in the exchange of water between the atmosphere and Earth’s surface. Whenever matter, including water, changes from one state to another, energy is either absorbed or released. The amount of energy released or absorbed (per gram) during a change of state is known as latent heat (from the Latin latens, “hidden”). In going from a more ordered state (liquid) to a less ordered state (gas), energy is absorbed (the energy required to break bonds between molecules). While going from a less ordered to a more ordered state, energy is released. Evaporation, the transformation of water from a liquid to a gaseous state, requires 2260 joules (J) of energy per gram of liquid water to be converted to water vapor (1 joule is the equivalent of 1 watt of power radiated or dissipated for 1 second). Condensation, the transformation of water vapor to a liquid state, releases an equivalent amount of energy. When air comes into contact with liquid water, water molecules are freely exchanged between the air and the water’s surface. When the evaporation rate equals the condensation rate, the air is said to be saturated. In the air, water vapor acts as an independent gas that has weight and exerts pressure. The amount of pressure that water vapor exerts independent of the pressure of dry air is called vapor pressure. Vapor pressure is typically defined in units of pascals (Pa). The water vapor content of air at saturation is called the saturation vapor pressure. The saturation vapor pressure, also known as the water vapor capacity of air, cannot be exceeded. If the vapor pressure exceeds the capacity, condensation occurs and reduces the vapor pressure. Saturation vapor pressure varies with temperature, increasing as air temperature increases (Figure 2.15). Having a greater quantity of thermal energy to support evaporation, warm air has a greater capacity for water vapor than does cold air. Interpreting Ecological Data Q1. Assume that the actual (current) water vapor pressure remains the same over the course of the day and that the current air temperature of 25°C in the above graph represents the air temperature at noon (12:00 p.m.). How would you expect the relative humidity to change from noon to 5:00 p.m.? Why? Q2. What is the approximate relative humidity at 35°C? (Assume that actual water vapor pressure remains the same as in the above figure, 2 kilopascals [kPa].) The amount of water in a given volume of air is its absolute humidity. A more familiar measure of the water content of the air is relative humidity, or the amount of water vapor in the air expressed as a percentage of the saturation vapor pressure. At saturation vapor pressure, the relative humidity is 100 percent. If air cools while the actual moisture content (water vapor pressure) remains constant, then relative humidity increases as the value of saturation vapor pressure declines. If the air cools to a point where the actual vapor pressure is equal to the saturation vapor pressure, moisture in the air will condense. This is what occurs when a warm parcel of air at the surface becomes buoyant and rises. As it rises, it cools, and as it cools, the relative humidity increases. When the relative humidity reaches 100 percent, water vapor condenses and forms clouds. As soon as particles of water or ice in the air become too heavy to remain suspended, precipitation falls. For a given water content of a parcel of air (vapor pressure), the temperature at which saturation vapor pressure is achieved (relative humidity is 100 percent) is called the dew point temperature. Think about finding dew or frost on a cool fall morning. As nightfall approaches, temperatures drop and relative humidity rises. If cool night air temperatures reach the dew point, water condenses and dew forms, lowering the amount of water in the air. As the sun rises, air temperature warms and the water vapor capacity (saturation vapor pressure) increases. As a result, the dew evaporates, increasing vapor pressure in the air. 2.6 Precipitation Has a Distinctive Global Pattern By bringing together patterns of temperature, winds, and ocean currents, we are ready to understand the global pattern of precipitation. Precipitation is not evenly distributed across Earth (Figure 2.16). At first the global map of annual precipitation in Figure 2.16 may seem to have no discernible pattern or regularity. But if we examine the simpler pattern of variation in average rainfall with latitude (Figure 2.17), a general pattern emerges. Precipitation is highest in the region of the equator, declining as one moves north and south. The decline, however, is not continuous. Two troughs occur in the midlatitudes interrupting the general patterns of decline in precipitation from the equator toward the poles. The sequence of peaks and troughs seen in Figure 2.17 corresponds to the pattern of rising and falling air masses associated with the belts of prevailing winds presented in Figure 2.12. As the warm trade winds move across the tropical oceans, they gather moisture. Near the equator, the northeasterly trade winds meet the southeasterly trade winds. This narrow region where the trade winds meet is the ITCZ, characterized by high amounts of precipitation. Where the two air masses meet, air piles up, and the warm humid air rises and cools. When the dew point is reached, clouds form, and precipitation falls as rain. This pattern accounts for high precipitation in the tropical regions of eastern Asia, Africa, and South and Central America (see Figure 2.16). Having lost much of its moisture, the ascending air mass continues to cool as it splits and moves northward and southward. In the region of the subtropical high (approximately 30° N and S), where the cool air descends, two belts of dry climate encircle the globe (the two troughs at the midlatitudes seen in Figure 2.17). The descending air warms. Because the saturation vapor pressure rises, it draws water from the surface through evaporation, causing arid conditions. In these belts, the world’s major deserts have formed (see Chapter 23). As the air masses continue to move north and south, they once again draw moisture from the surface, but to a lesser degree because of the cooler surface conditions. Moving poleward, they encounter cold air masses originating at the poles (approximately 60° N and S). Where the surface air masses converge and rise, the ascending air mass cools and precipitation occurs (seen as the two smaller peaks in precipitation between 50° and 60° N and S in Figure 2.17). From this point on to the poles, the cold temperature and associated low-saturation vapor pressure function to restrict precipitation. One other pattern is worth noting in Figure 2.17. In general, rainfall is greater in the Southern Hemisphere than in the Northern Hemisphere (note the southern shift in the rainfall peak associated with the ITCZ). This is because the oceans cover a greater proportion of the Southern Hemisphere, and water evaporates more readily from the water’s surface than from the soil and vegetation. Missing from our discussion thus far is the temporal variation of precipitation over Earth. The temporal variation is directly linked to the seasonal changes in the surface radiation balance of Earth and its effect on the movement of global pressure systems and air masses. This is illustrated in seasonal movement north and south of the ITCZ, which follows the apparent migration of the direct rays of the Sun (Figure 2.18). The ITCZ is not stationary but tends to migrate toward regions of the globe with the warmest surface temperature. Although tropical regions around the equator are always exposed to warm temperatures, the Sun is directly over the geographical equator only twice a year, at the spring and fall equinoxes. At the northern summer solstice, the Sun is directly over the Tropic of Cancer; at the winter solstice (which is summer in the Southern Hemisphere), the Sun is directly over the Tropic of Capricorn. As a result, the ITCZ moves poleward and invades the subtropical highs in northern summer; in the winter it moves southward, leaving clear, dry weather behind. As the ITCZ migrates southward, it brings rain to the southern summer. Thus, as the ITCZ shifts north and south, it brings on the wet and dry seasons in the tropics (Figure 2.19). 2.7 Proximity to the Coastline Influences Climate At the continental scale, an important influence on climate is the relationship between land and water. Land surfaces heat and cool more rapidly than water as a result of differences in their specific heat. Specific heat is the amount of thermal energy necessary to raise the temperature of one gram of a substance by 1°C. The specific heat of water is much higher than that of land or air. It takes approximately four times the amount of thermal energy to raise the temperature of water by 1°C than land or air. As a result, land areas farther from the coast (or other large bodies of water) experience a greater seasonal variation in temperature than do coastal areas (Figure 2.20). This pattern is referred to as continentality. Annual differences of as much as 100°C (from 50°C to –50°C) have been recorded in some locations. The converse effect occurs in coastal regions. These locations have smaller temperature ranges as a result of what is called a maritime influence. Summer and winter extremes are moderated by the movement onshore of prevailing westerly wind systems from the ocean. Ocean currents minimize seasonal variations in the surface temperature of the water. The moderated water temperature serves to moderate temperature changes in the air mass above the surface. Proximity to large water bodies also tends to have a positive influence on precipitation levels. The interior of continents generally experience less precipitation than the coastal regions do. As air masses move inland from the coast, water vapor lost from the atmosphere through precipitation is not recharged (from surface evaporation) as readily as it is over the open waters of the ocean (note the gradients of precipitation from the coast to the interiors of North America and Europe/Asia in Figure 2.16). There are, however, notable exceptions to this rule, including the dry coast of southern California and the Arctic coastline of Alaska. 2.8 Topography Influences Regional and Local Patterns of Climate Mountainous topography influences local and regional patterns of climate. Most obvious is the relationship between elevation and temperature. In the lower regions of the atmosphere (up to altitudes of approximately 12 km), temperature decreases with altitude at a fairly uniform rate because of declining air density and pressure. In addition, the atmosphere is warmed by conduction (transfer of heat through direct contact) from Earth’s surface. So temperature declines with increasing distance from the conductive source (i.e., the surface). The rate of decline in temperature with altitude is called the lapse rate. So for the same latitude or proximity to the coast, locales at higher elevation will have consistently lower temperatures than those of lower elevation. Mountains also influence patterns of precipitation. As an air mass reaches a mountain, it ascends, cools, relative humidity rises (because of lower saturation vapor pressure). When the temperature cools to the dew point temperature, precipitation occurs at the upper altitudes of the windward side. As the now cool, dry air descends the leeward side, it warms again and relative humidity declines. As a result, the windward side of a mountain supports denser, more vigorous vegetation and different species of plants and associated animals than does the leeward side, where in some areas dry, desert-like conditions exist. This phenomenon is called a rain shadow (Figure 2.21). Thus, in North America, the westerly winds that blow over the Sierra Nevada and the Rocky Mountains, dropping their moisture on west-facing slopes, support vigorous forest growth. By contrast, the eastern slopes exhibit semi-desert or desert conditions. Some of the most pronounced effects of this same phenomenon occur in the Hawaiian Islands. There, plant cover ranges from scrubby vegetation on the leeward side of an island to moist, forested slopes on the windward side (Figure 2.22). 2.9 Irregular Variations in Climate Occur at the Regional Scale The patterns of temporal variation in climate that we have discussed thus far occur at regular and predictable intervals: seasonal changes in temperature with the rotation of Earth around the Sun, and migration of the ITCZ with the resultant seasonality of rainfall in the tropics and monsoons in Southeast Asia. Not all features of the climate system, however, occur so regularly. Earth’s climate system is characterized by variability at both the regional and global scales. The Little Ice Age, a period of cooling that lasted from approximately the mid-14th to the mid-19th century, brought bitterly cold winters to many parts of the Northern Hemisphere, affecting agriculture, health, politics, economics, emigration, and even art and literature. In the mid-17th century, glaciers in the Swiss Alps advanced, gradually engulfing farms and crushing entire villages. In 1780, New York Harbor froze, allowing people to walk from Manhattan to Staten Island. In fact, the image of a white Christmas evoked by Charles Dickens and the New England poets of the 18th and 19th centuries is largely a product of the cold and snowy winters of the Little Ice Age. But the climate has since warmed to the point that a white Christmas in these regions is becoming an anomaly. The Great Plains region of central North America has undergone periods of drought dating back to the mid-Holocene period some 5000 to 8000 years ago, but the homesteaders of the early 20th century settled the Great Plains at a time of relatively wet summers. They assumed these moisture conditions were the norm, and they employed the agricultural methods they had used in the East. So they broke the prairie sod for crops, but the cycle of drought returned, and the prairie grasslands became a dust bowl (see Chapter 4, Ecological Issues & Applications). These examples reflect the variability in Earth’s climate systems, which operate on timescales ranging from decades to tens of thousands of years, driven by changes in the input of energy to Earth’s surface (see Section 2.1). Earth’s orbit is not permanent. Changes occur in the tilt of the axis and the shape of the yearly path about the Sun. These variations affect climate by altering the seasonal inputs of solar radiation. Occurring on a timescale of tens of thousands of years, these variations are associated with the glacial advances and retreats throughout Earth’s history (see Chapter 18). Variations in the level of solar radiation to Earth’s surface are also associated with sunspot activity—huge magnetic storms on the Sun. These storms are associated with strong solar emissions and occur in cycles, with the number and size reaching a maximum approximately every 11 years. Researchers have related sunspot activity, among other occurrences, to periods of drought and winter warming in the Northern Hemisphere. Interaction between two components of the climate system, the ocean and the atmosphere, are connected to some major climatic variations that occur at a regional scale. As far back as 1525, historic documents reveal that fishermen off the coast of Peru recorded periods of unusually warm water. The Peruvians referred to these as El Niño because they commonly appear at Christmastime, the season of the Christ Child (Spanish: El Niño). Now referred to by scientists as the El Niño–Southern Oscillation (ENSO), this phenomenon is a global event arising from large-scale interaction between the ocean and the atmosphere. The Southern Oscillation, a more recent discovery, refers to an oscillation in the surface pressure (atmospheric mass) between the southeastern tropical Pacific and the Australian-Indonesian regions. When the waters of the eastern Pacific are abnormally warm (an El Niño event), sea level pressure drops in the eastern Pacific and rises in the west. The reduction in the pressure gradient is accompanied by a weakening of the low-latitude easterly trades. Although scientists still do not completely understand the cause of the ENSO phenomenon, its mechanism has been well documented. Recall from Section 2.3 that the trade winds blow westward across the tropical Pacific (see Figure 2.12). As a consequence, the surface currents within the tropical oceans flow westward (see Figure 2.14), bringing cold, deeper waters to the surface off the coast of Peru in a process known as upwelling (see Section 3.8). This pattern of upwelling, together with the cold-water current flowing from south to north along the western coast of South America, results in this region of the ocean usually being colder than one would expect given its equatorial location (Figure 2.23). As the surface currents move westward the water warms, giving the water’s destination, the western Pacific, the warmest ocean surface on Earth. The warmer water of the western Pacific causes the moist maritime air to rise and cool, bringing abundant rainfall to the region (Figure 2.23; also see Figure 2.16). In contrast, the cooler waters of the eastern Pacific result in relatively dry conditions along the Peruvian coast. During an El Niño event, the trade winds slacken, reducing the westward flow of the surface currents (see Figure 2.23). The result is a reduced upwelling and a warming of the surface waters in the eastern Pacific. Rainfall follows the warm water eastward, with associated flooding in Peru and drought in Indonesia and Australia. This eastward displacement of the atmospheric heat source (latent heat associated with the evaporation of water; see Section 3.2) overlaying the warm surface waters results in large changes in global atmospheric circulation, in turn influencing weather in regions far removed from the tropical Pacific. At other times, the injection of cold water becomes more intense than usual, causing the surface of the eastern Pacific to cool. This variation is referred to as La Niña (Figure 2.24). It results in droughts in South America and heavy rainfall, even floods, in eastern Australia. 2.10 Most Organisms Live in Microclimates Most organisms live in local conditions that do not match the general climate profile of the larger region surrounding them. For example, today’s weather report may state that the temperature is 28°C and the sky is clear. However, your weather forecaster is painting only a general picture. Actual conditions of specific environments will be quite different depending on whether they are underground versus on the surface, beneath vegetation or on exposed soil, or on mountain slopes or at the seashore. Light, heat, moisture, and air movement all vary greatly from one part of the landscape to another, influencing the transfer of heat energy and creating a wide range of localized climates. These microclimates define the conditions organisms live in. On a sunny but chilly day in early spring, flies may be attracted to sap oozing from the stump of a maple tree. The flies are active on the stump despite the near-freezing air temperature because, during the day, the surface of the stump absorbs solar radiation, heating a thin layer of air above the surface. On a still day, the air heated by the tree stump remains close to the surface, and temperatures decrease sharply above and below this layer. A similar phenomenon occurs when the frozen surface of the ground absorbs solar radiation and thaws. On a sunny, late winter day, the ground is muddy even though the air is cold. By altering soil temperatures, moisture, wind movement, and evaporation, vegetation moderates microclimates, especially areas near the ground. For example, areas shaded by plants have lower temperatures at ground level than do places exposed to the Sun. On fair summer days in locations 25 millimeters (mm; 1 inch) aboveground, dense forest cover can reduce the daily range of temperatures by 7°C to 12°C below the soil temperature in bare fields. Under the shelter of heavy grass and low plant cover, the air at ground level is completely calm. This calm is an outstanding feature of microclimates within dense vegetation at Earth’s surface. It influences both temperature and humidity, creating a favorable environment for insects and other ground-dwelling animals. Topography, particularly aspect (the direction that a slope faces), influences the local climatic conditions. In the Northern Hemisphere, south-facing slopes receive the most solar energy, whereas north-facing slopes receive the least (Figure 2.25). At other slope positions, energy received varies between these extremes, depending on their compass direction. Different exposure to solar radiation at south- and north-facing sites has a marked effect on the amount of moisture and heat present. Microclimate conditions range from warm, dry, variable conditions on the south-facing slope to cool, moist, more uniform conditions on the north-facing slope. Because high temperatures and associated high rates of evaporation draw moisture from soil and plants, the evaporation rate at south-facing slopes is often 50 percent higher, the average temperature is higher, and soil moisture is lower. Conditions are driest on the tops of south-facing slopes, where air movement is greatest, and dampest at the bottoms of north-facing slopes. The same microclimatic conditions occur on a smaller scale on north- and south-facing slopes of large ant hills, mounds of soil, dunes, and small ground ridges in otherwise flat terrain, as well as on the north- and south-facing sides of buildings, trees, and logs. The south-facing sides of buildings are always warmer and drier than the north-facing sides—a consideration for landscape planners, horticulturists, and gardeners. North sides of tree trunks are cooler and moister than south sides, as reflected by more vigorous growth of moss on the north sides. In winter, the temperature of the north-facing side of a tree may be below freezing while the south side, heated by the Sun, is warm. This temperature difference may cause frost cracks in the bark as sap, thawed by day, freezes at night. Bark beetles and other wood-dwelling insects that seek cool, moist areas for laying their eggs prefer north-facing locations. Flowers on the south side of tree crowns often bloom sooner than those on the north side. Microclimatic extremes also occur in depressions in the ground and on the concave surfaces of valleys, where the air is protected from the wind. Heated by sunlight during the day and cooled by terrestrial vegetation at night, this air often becomes stagnant. As a result, these sheltered sites experience lower nighttime temperatures (especially in winter), higher daytime temperatures (especially in summer), and higher relative humidity. If the temperature drops low enough, frost pockets form in these depressions. The microclimates of the frost pockets often display the same phenomenon, supporting different kinds of plant life than found on surrounding higher ground. Interpreting Ecological Data Q1. Which of the two slope positions (north- or south-facing) has the higher maximum recorded temperatures (mid-afternoon)? Q2. How does vegetation cover (forested vs. exposed slope) influence surface temperatures? Although the global and regional patterns of climate discussed constrain the large-scale distribution and abundance of plants and animals, the localized patterns of microclimate define the actual environmental conditions sensed by the individual organism. This localized microclimate thus determines the distribution and activities of organisms in a particular region. Ecological Issues & Applications Rising Atmospheric Concentrations of Greenhouse Gases Are Altering Earth’s Climate Since the middle of the 19th century, direct measurements of surface temperature have been made at widespread locations around the world. These direct measures from instruments such as thermometers are referred to as the instrumental record. Besides these measurements made at the land surface, observations of sea surface temperatures have been made from ships since the mid-19th century. Since the late 1970s, both a network of instrumented buoys and Earth-observing satellites have been providing a continuous record of global observations for a wide variety of climate variables, supplementing the previous land- and ship-based instrumental records. What these various sources of data on the land and sea surface temperatures of our planet indicate is that Earth has been warming over the past 150 years (Figure 2.26). Since the early 20th century, the global average surface temperature has increased by 0.74°C (±0.2°C). In addition, the 10 warmest years in the instrumental record since 1850 are, in descending order, 2010, 2005, 1998, 2003, 2013, 2002, 2006, 2009, 2007, and 2004. Analyses also indicate that global ocean heat content has increased significantly since the late 1950s. More than half of the increase in heat content has occurred in the upper 300 meters of the ocean; in this layer the temperature has increased at a rate of about 0.04°C per decade. Additional data examining trends on humidity, sea-ice extent, and snow cover likewise indicate a pattern of warming over the past century. What is the cause of this warming? The scientific consensus is that the warming is in large part a result of rising atmospheric concentrations of greenhouse gases. According to the most recent report of the Intergovernmental Panel on Climate Change (Report of Working Group I, 2013): Warming of the climate system is unequivocal, as is now evident from observations of increases in global average air and ocean temperatures . . . . Most of the observed increase in global average temperatures since the mid-20th century is very likely due to the observed increase in anthropogenic greenhouse gas concentrations. Although human activities have increased the atmospheric concentration of a variety of greenhouse gases (e.g., methane [CH4], nitrous oxide [N2O]), the major concern is focused on carbon dioxide (CO2). The atmospheric concentration of CO2 has increased by more than 30 percent over the past 100 years. The evidence for this rise comes primarily from continuous observations of atmospheric CO2 started in 1958 at Mauna Loa, Hawaii, by Charles Keeling (Figure 2.27) and from parallel records around the world. Evidence before the direct observations of 1958 comes from various sources, including the analysis of air bubbles trapped in the ice of glaciers in Greenland and Antarctica. In reconstructing atmospheric CO2 concentrations over the past 300 years, we see values that fluctuate between 280 and 290 parts per million (ppm) until the mid-1800s (see Figure 2.27). After the onset of the Industrial Revolution, the value increased steadily, rising exponentially by the mid-19th century onward. The change reflects the combustion of fossil fuels (coal, oil, and gas) as an energy source for industrialized nations (Figure 2.28a), as well as the increased clearing and burning of forests (primarily in the tropical regions; see Figure 2.28b). Although there is an obvious correlation between rising atmospheric concentrations of CO2 (and other greenhouse gases) and the observed increases in global temperature, what makes the scientific community so confident that the observed rise in global temperatures is a result of the greenhouse effect? One important factor is the actual pattern of warming itself. Recall from our discussion of the Earth’s radiation balance, that surface temperature at any location or time reflects the net radiation balance, that is, the difference between incoming shortwave radiation and outgoing longwave radiation (Section 2.1). If incoming shortwave radiation exceeds outgoing longwave radiation, surface temperatures rise. Conversely, if outgoing longwave radiation exceeds incoming shortwave radiation, temperatures decline. It is this imbalance that accounts for the decline in mean annual temperatures with increasing latitude from the tropics (net radiation surplus) to the poles (net radiation deficit; see Figure 2.6). Likewise, it is the shift from surplus to deficit that results in the decline in surface temperatures from day to night (diurnal cycle) and from summer to winter (seasonal cycle). Since the influence of greenhouse gases on the radiation balance works through the absorption of outgoing longwave radiation, which is then emitted downward toward the surface instead, the net effect reduces cooling, that is, keeps the surface temperature warmer than it would otherwise be if the longwave radiation were lost to space. It therefore follows that the greater proportional warming from rising levels of greenhouse gases would occur in those places (i.e., polar) and times (i.e., winter and night) when and where temperatures are generally declining as a result of negative net radiation balance. An analysis of the patterns of warming over the past 50 years is in general agreement with this expectation. The increase in global mean surface temperature illustrated in Figure 2.27 has not been the same at every location. The global map presented in Figure 2.29a shows the geographic patterns of surface temperature changes over the period from 1955 to 2005. Note that the greatest warming has occurred in the polar regions, particularly the Arctic (North America and Eurasia between 40 and 70° N). Although Earth’s average temperature has risen 0.74°C during the 20th century, the Arctic is warming twice as fast as other parts of the world. In Alaska (U.S.) average temperatures have increased 3.0°C between 1970 and 2000. The warmer temperatures have caused other changes in the Arctic region such as melting of sea ice and continental ice sheets (Greenland). The reduction in ice cover potentially exacerbates the problem by reducing surface albedo and increasing the absorption of incoming shortwave radiation. In the Southern hemisphere, the Antarctic Peninsula has also undergone a great warming—five times the global average. Interpreting Ecological Data Q1. Based on the data provided in (b), which latitudes exhibit the greatest seasonal variations in surface temperature (Ts) change? Q2. What accounts for the fact that the period of Jun–Aug in the arctic region (north of 60° N) shows the least warming, while the same period corresponds to the maximum temperature change in the Antarctic (south of 60° S)? The changes in mean surface temperature presented in Figure 2.29a have been partitioned by season (December–February, March–May, June–August, and September–December.) in Figure 2.29b. The greatest observed warming over the last half century has occurred during the winter months. In effect, this represents a reduction of the normal pattern of cooling that occurs during the winter months as a result of the deficit in net radiation (deficit is reduced by increased absorption of outgoing longwave radiation). This pattern of winter warming becomes more apparent when the seasonal data are analyzed by latitude (see Figure 2.29b). The net result of winter warming is a reduction in the seasonal variations in temperature (differences between the warmest and coldest months). Analyses of daily maximum and minimum land-surface temperatures from 1950 to 2000 show a decrease in the diurnal temperature range. On average, minimum temperatures are increasing at about twice the rate of maximum temperatures (0.2°C versus 0.1°C per decade). In other words, nighttime temperatures (minimum) have increased more than daytime temperatures (maximum) over this period. These patterns of increasing surface temperatures over the past century have a major influence on the functioning of ecological systems, arranging the distribution of plant and animal species, the structure of communities, and the patterns of ecosystem productivity and decomposition. We will explore a variety of these issues in the Ecological Issues & Applications sections of the chapters that follow, and examine in more detail the current and future implication of global climate change in Chapter 27. Summary Net Radiation 2.1 Earth intercepts solar energy in the form of shortwave radiation, some of which is reflected back into space. Earth emits energy back into space in the form of longwave radiation, a portion of which is absorbed by gases in the atmosphere and radiated back to the surface. The difference between incoming shortwave and outgoing longwave radiation is the net radiation. Surface temperatures are a function of net radiation. Seasonal Variation 2.2 The amount of solar radiation intercepted by Earth varies markedly with latitude. Tropical regions near the equator receive the greatest amount of solar radiation, and high latitudes receive the least. Because Earth tilts on its axis, parts of Earth encounter seasonal differences in solar radiation. These differences give rise to seasonal variations in net radiation and temperature. There is a global gradient in mean annual temperature; it is warmest in the tropics and declines toward the poles. Atmospheric Circulation 2.3 From the equator to the midlatitudes there is an annual surplus of net radiation, and there is a deficit from the midlatitudes to the poles. This latitudinal gradient of net radiation gives rise to global patterns of atmospheric circulation. The spin of Earth on its axis deflects air and water currents to the right in the Northern Hemisphere and to the left in the Southern Hemisphere. Three cells of global air flow occur in each hemisphere. Ocean Currents 2.4 The global pattern of winds and the Coriolis effect cause major patterns of ocean currents. Each ocean is dominated by great circular water motions, or gyres. These gyres move clockwise in the Northern Hemisphere and counterclockwise in the Southern Hemisphere. Atmospheric Moisture 2.5 Atmospheric moisture is measured in terms of relative humidity. The maximum amount of moisture the air can hold at any given temperature is called the saturation vapor pressure, which increases with temperature. Relative humidity is the amount of water in the air, expressed as a percentage of the maximum amount the air could hold at a given temperature. Precipitation 2.6 Wind, temperature, and ocean currents produce global patterns of precipitation. They account for regions of high precipitation in the tropics and belts of dry climate at approximately 30° N and S latitude. Continentality 2.7 Land surfaces heat and cool more rapidly than water; as a result, land areas farther from the coast experience a greater seasonal variation in temperature than do coastal areas. The interiors of continents generally receive less precipitation than the coastal regions do. Topography 2.8 Temperature declines with altitude, so locations at higher elevations will have consistently lower temperatures that those of lower elevations. Mountainous topography influences local and regional patterns of precipitation. As an air mass reaches a mountain, it ascends, cools, becomes saturated with water vapor, and releases much of its moisture at upper altitudes of the windward side. Irregular Variation 2.9 Not all temporal variation in regional climate occurs at a regular interval. Irregular variations in the trade winds give rise to periods of unusually warm waters off the coast of western South America. Referred to by scientists as El Niño; this phenomenon is a global event arising from large-scale interaction between the ocean and the atmosphere. Microclimates 2.10 The actual climatic conditions that organisms live in vary considerably within one climate. These local variations, or microclimates, reflect topography, vegetative cover, exposure, and other factors on every scale. Angles of solar radiation cause marked differences between north- and south-facing slopes, whether on mountains, sand dunes, or ant mounds. Climate Warming Ecological Issues & Applications Over the past century the average surface temperature of the planet has been rising. The rise in surface temperature is related to increasing atmospheric concentrations of greenhouse gases caused by the burning of fossil fuels and clearing and burning of forests.
How to Measure and Use Leaf Area Index
CHAPTER 3 Smith, T. M., & Smith, R. L. (2015). Elements of Ecology (9th ed.). Boston, MA: Pearson. 3.1 Water Cycles between Earth and the Atmosphere All marine and freshwater aquatic environments are linked, either directly or indirectly, as components of the water cycle (also referred to as the hydrologic cycle; Figure 3.1)—the process by which water travels in a sequence from the air to Earth and returns to the atmosphere. Solar radiation, which heats Earth’s atmosphere and provides energy for the evaporation of water, is the driving force behind the water cycle (see Chapter 2). Precipitation sets the water cycle in motion. Water vapor, circulating in the atmosphere, eventually falls in some form of precipitation. Some of the water falls directly on the soil and bodies of water. Some is intercepted by vegetation, dead organic matter on the ground, and urban structures and streets in a process known as interception. Because of interception, which can be considerable, various amounts of water never infiltrate the ground but evaporate directly back to the atmosphere. Precipitation that reaches the soil moves into the ground by infiltration. The rate of infiltration depends on the type of soil, slope, vegetation, and intensity of the precipitation (see Section 4.8). During heavy rains when the soil is saturated, excess water flows across the surface of the ground as surface runoff or overland flow. At places, it concentrates into depressions and gullies, and the flow changes from sheet to channelized flow—a process that can be observed on city streets as water moves across the pavement into gutters. Because of low infiltration, runoff from urban areas might be as much as 85 percent of the precipitation. Some water entering the soil seeps down to an impervious layer of clay or rock to collect as groundwater (see Figure 3.1). From there, water finds its way into springs and streams. Streams coalesce into rivers as they follow the topography of the landscape. In basins and floodplains, lakes and wetlands form. Rivers eventually flow to the coast, forming the transition from freshwater to marine environments. Water remaining on the surface of the ground, in the upper layers of the soil, and collected on the surface of vegetation—as well as water in the surface layers of streams, lakes, and oceans—returns to the atmosphere by evaporation. The rate of evaporation is governed by how much water vapor is in the air relative to the saturation vapor pressure (relative humidity; see Section 2.5). Plants cause additional water loss from the soil. Through their roots, they take in water from the soil and lose it through their leaves and other organs in a process called transpiration. Transpiration is the evaporation of water from internal surfaces of leaves, stems, and other living parts (see Chapter 6). The total amount of evaporating water from the surfaces of the ground and vegetation (surface evaporation plus transpiration) is called evapotranspiration. Figure 3.2 is a diagram of the global water cycle showing the various reservoirs (bodies of water) and fluxes (exchanges between reservoirs). The total volume of water on Earth is approximately 1.4 billion cubic kilometers (km3) of which more than 97 percent resides in the oceans. Another 2 percent of the total is found in the polar ice caps and glaciers, and the third-largest active reservoir is groundwater (0.3 percent). Over the oceans, evaporation exceeds precipitation by some 40,000 km3. A significant proportion of the water evaporated from the oceans is transported by winds over the land surface in the form of water vapor, where it is deposited as precipitation. Of the 111,000 km3 of water that falls as precipitation on the land surface, only some 71,000 km3 is returned to the atmosphere as evapotranspiration. The remaining 40,000 km3 is carried as runoff by rivers and eventually returns to the oceans. This amount balances the net loss of water from the oceans to the atmosphere through evaporation that is eventually deposited on the continents (land surface) as precipitation (see Figure 3.2). The relatively small size of the atmospheric reservoir (only 13 km3) does not reflect its importance in the global water cycle. In Figure 3.2, note the large fluxes between the atmosphere, the oceans, and the land surface relative to the amount of water residing in the atmosphere at any given time (e.g., the size of atmospheric reservoir). The importance of the atmosphere in the global water cycle is better reflected by the turnover time of this reservoir. The turnover time is calculated by dividing the size of the reservoir by the rate of output (flux out). For example, the turnover time for the ocean is the size of the reservoir (1.37 × 106 km3) divided by the rate of evaporation (425 km3 per year) or more than 3000 years. In contrast, the turnover time of the atmospheric reservoir is approximately 0.024 year. That is to say, the entire water content of the atmosphere is replaced on average every nine days. 3.2 Water Has Important Physical Properties The physical arrangement of its component molecules makes water a unique substance. A molecule of water consists of two atoms of hydrogen (H) joined to one atom of oxygen (O), represented by the chemical symbol H2O. The H atoms are bonded to the O atom asymmetrically, such that the two H atoms are at one end of the molecule and the O atom is at the other (Figure 3.3a). The bonding between the two hydrogen atoms and the oxygen atom is via shared electrons (called a covalent bond), so that each H atom shares a single electron with the oxygen. The shared hydrogen atoms are closer to the oxygen atom than they are to each other. As a result, the side of the water molecule where the H atoms are located has a positive charge, and the opposite side where the oxygen atom is located has a negative charge, thus polarizing the water molecule (termed a polar covalent bond; Figure 3.3b). Because of its polarity, each water molecule becomes weakly bonded with its neighboring molecules (Figure 3.3c). The positive (hydrogen) end of one molecule attracts the negative (oxygen) end of the other. The angle between the hydrogen atoms encourages an open, tetrahedral arrangement of water molecules. This situation, wherein hydrogen atoms act as connecting links between water molecules, is called hydrogen bonding. The simultaneous bonding of a hydrogen atom to the oxygen atoms of two different water molecules gives rise to a lattice arrangement of molecules (Figure 3.3d). These bonds, however, are weak in comparison to the bond between the hydrogen and oxygen atoms. As a result, they are easily broken and reformed. Water has some unique properties related to its hydrogen bonds. One property is high specific heat—the number of calories necessary to raise the temperature of 1 gram of water 1 degree Celsius. The specific heat of water is defined as a value of 1, and other substances are given a value relative to that of water. Water can store tremendous quantities of heat energy with a small rise in temperature. As a result, great quantities of heat must be absorbed before the temperature of natural waters, such as ponds, lakes, and seas, rises just 1°C. These waters warm up slowly in spring and cool off just as slowly in the fall. This process prevents the wide seasonal fluctuations in the temperature of aquatic habitats so characteristic of air temperatures and moderates the temperatures of local and worldwide environments (see Section 2.7). The high specific heat of water is also important in the thermal regulation of organisms. Because 75–95 percent of the weight of all living cells is water, temperature variation is also moderated relative to changes in ambient temperature. As a result of the high specific heat of water, large quantities of heat energy are required for it to change its state between solid (ice), liquid, and gaseous (water vapor) phases. Collectively, the energy released or absorbed in transforming water from one state to another is called latent heat (see Section 2.5). Removing only 1 calorie (4.184 joules [J]) of heat energy will lower the temperature of a gram of water from 2°C to 1°C, but approximately 80 times as much heat energy (80 calories per gram) must be removed to convert that same quantity of water at 1°C to ice (water’s freezing point of 0°C). Likewise, it takes 536 calories to overcome the attraction between molecules and convert 1 gram (g) of water at 100°C into vapor, the same amount of heat needed to raise 536 g of water 1°C. The lattice arrangement of molecules gives water a peculiar density–temperature relationship. Most liquids become denser as they are cooled. If cooled to their freezing temperature, they become solid, and the solid phase is denser than the liquid. This description is not true for water. Pure water becomes denser as it is cooled until it reaches 4°C ( Figure 3.4 ). Cooling below this temperature results in a decrease in density. When 0°C is reached, freezing occurs and the lattice structure is complete—each oxygen atom is connected to four other oxygen atoms by means of hydrogen atoms. The result is a lattice with large, open spaces and therefore decreased density (see Figure 3.3e). When frozen, water molecules occupy more space than they do in liquid form. Because of its reduced density, ice is lighter than water and floats on it. This property is crucial to life in aquatic environments. The ice on the surface of water bodies insulates the waters below, helping to keep larger bodies of water from freezing solid during the winter months. Because of hydrogen bonding, water molecules tend to stick firmly to one another, resisting external forces that would break their bonds. This property is called cohesion. In a body of water, these forces of attraction are the same on all sides. At the water’s surface, however, conditions are different. Below the surface, molecules of water are strongly attracted to one another. Above the surface is the much weaker attraction between water molecules and air. Therefore, molecules on the surface are drawn downward, resulting in a surface that is taut like an inflated balloon. This condition, called surface tension, is important in the lives of aquatic organisms. For example, the surface of water is able to support small objects and animals, such as the water striders (Gerridae spp.) and water spiders (Dolomedes spp.) that run across a pond’s surface ( Figure 3.5 ). To other small organisms, surface tension is a barrier, whether they wish to penetrate the water below or escape into the air above. For some, the surface tension is too great to break; for others, it is a trap to avoid while skimming the surface to feed or to lay eggs. If caught in the surface tension, a small insect may flounder on the surface. The nymphs of mayflies (Ephemeroptera spp.) and caddis flies (Trichoptera spp.) that live in the water and transform into winged adults are hampered by surface tension when trying to emerge from the water. While slowed down at the surface, these insects become easy prey for fish. Cohesion is also responsible for the viscosity of water. Viscosity is the property of a material that measures the force necessary to separate the molecules and allow an object to pass through the liquid. Viscosity is the source of frictional resistance to objects moving through water. This frictional resistance of water is 100 times greater than that of air. The streamlined body shape of many aquatic organisms, for example most fish and marine mammals, helps to reduce frictional resistance. Replacement of water in the space left behind by the moving animal increases drag on the body. An animal streamlined in reverse, with a short, rounded front and a rapidly tapering body, meets the least water resistance. The perfect example of such streamlining is the sperm whale (Physeter catodon; Figure 3.6 ). Water’s high viscosity relative to that of air is largely the result of its greater density. The density of water is about 860 times greater than that of air (pure water has a density of 1000 kilograms per cubic meter [kg/m3]). Although the resulting viscosity of water limits the mobility of aquatic organisms, it also benefits them. If a body submerged in water weighs less than the water it displaces, it is subjected to an upward force called buoyancy. Because most aquatic organisms (plants and animals) are close to neutral buoyancy (their density is similar to that of water), they do not require structural material such as skeletons or cellulose to hold their bodies erect against the force of gravity. Similarly, when moving on land, terrestrial animals must raise their mass against the force of gravity with each step they take. Such movement requires significantly more energy than swimming movements do for aquatic organisms. But water’s greater density can profoundly affect the metabolism of marine organisms inhabiting the deeper waters of the ocean. Because of its greater density, water also undergoes greater changes in pressure with depth than does air. At sea level, the weight of the vertical column of air from the top of the atmosphere to the sea surface is 1 kilogram per square centimeter (kg/cm2) or 1 atmosphere (atm). In contrast, pressure increases 1 atm for each 10 m in depth. Because the deep ocean varies in depth from a few hundred meters down to the deep trenches at more than 10,000 m, the range of pressure at the ocean bottom is from 20 atm to more than 1000 atm. Recent research has shown that both proteins and biological membranes are strongly affected by pressure, and animals living in the deep ocean have evolved adaptations that allow these biochemical systems to function under conditions of extreme pressure. 3.3 Light Varies with Depth in Aquatic Environments When light strikes the surface of water, a certain amount is reflected back to the atmosphere. The amount of light reflected from the surface depends on the angle at which the light strikes the surface. The lower the angle, the larger the amount of light reflected. As a result, the amount of light reflected from the water surface will vary both diurnally and seasonally between the equator and the poles (see Section 2.1 and Figure 2.5 for a complete discussion). The amount of light entering the water surface is further reduced by two additional processes. First, suspended particles, both alive and dead, intercept the light and either absorb or scatter it. The scattering of light increases the length of its path through the water and results in further attenuation. Second, water itself absorbs light (Figure 3.7). Moreover, water absorbs some wavelengths more than others. First to be absorbed are visible red light and infrared radiation in wavelengths greater than 750 nanometers (nm). This absorption reduces solar energy by half. Next, in clear water, yellow disappears, followed by green and violet, leaving only blue wavelengths to penetrate deeper water. A fraction of blue light is lost with increasing depth. In the clearest seawater, only about 10 percent of blue light reaches to more than 100 m in depth. These changes in the quantity and quality of light have important implications for life in aquatic environments, both by directly influencing the quantity and distribution of productivity and by indirectly influencing the vertical profile of temperature with water depth (see Section 20.4 and Chapter 24). The lack of light in deeper waters of the oceans has resulted in various adaptations. Organisms of the deeper ocean (200–1000 m deep) are typically silvery gray or deep black, and organisms living in even deeper waters (below 1000 m) often lack pigment. Another adaptation is large eyes, which give these organisms maximum light-gathering ability. Many organisms have adapted organs that produce light through chemical reactions referred to as bioluminescence (see Section 24.10). Interpreting Ecological Data Q1. As you dive down in depth from the surface, which wavelength of light is the first to disappear? At approximately what depth would this occur? Q2. Is it the shorter or longer wavelengths of visible light that penetrate the deepest into the water column? (Refer to Figure 2.1.) 3.4 Temperature Varies with Water Depth Surface temperatures reflect the balance of incoming and outgoing radiation (see Section 2.1). As solar radiation is absorbed in the vertical water column, the temperature profile with depth might be expected to resemble the vertical profile of light shown in Figure 3.7—that is, decreasing exponentially with depth. However, the physical characteristic of water density plays an important role in modifying this pattern (see Section 3.2, Figure 3.4). As sunlight is absorbed in the surface waters, it heats up (Figure 3.8). Winds and surface waves mix the surface waters, distributing the heat vertically. Warm surface waters move downward, whereas the cooler waters below move up to the surface. As a result of this vertical mixing, heat is transported from the surface downward and the decline in water temperature with depth lags the decline in solar radiation. Below this mixed layer, however, temperatures drop rapidly. The region of the vertical depth profile where the temperature declines most rapidly is called the thermocline. The depth of the thermocline will depend on the input of solar radiation to the surface waters and on the degree of vertical mixing (wind speed and wave action). Below the thermocline, water temperatures continue to fall with depth but at a much slower rate. The result is a distinct pattern of temperature zonation with depth. The difference in temperature between the warm, well-mixed surface layer and the cooler waters below the thermocline causes a distinctive difference in water density in these two vertical zones. The thermocline is located between an upper layer of warm, lighter (less dense) water called the epilimnion and a deeper layer of cold, denser water called the hypolimnion (see Figure 3.8; also see Section 21.10 and Figure 21.23). The density change at the thermocline acts as a physical barrier that prevents mixing of the upper (epilimnion) and lower (hypolimnion) layers. Just as seasonal variation in the input of solar radiation to Earth’s surface results in seasonal changes in surface temperatures (see Section 2.2), seasonal changes in the input of solar radiation to the water surface give rise to seasonal changes in the vertical profile of temperature in aquatic environments (Figure 3.9). Because of the relatively constant input of solar radiation to the water surface throughout the year, the thermocline is a permanent feature of tropical waters. In the waters of the temperate zone, a distinct thermocline exists during the summer months. By fall, conditions begin to change, and a turnabout takes place. Air temperatures and sunlight decrease, and the surface water of the epilimnion starts to cool. As it does, the water becomes denser and sinks, displacing the warmer water below to the surface, where it cools in turn. As the difference in water density between the epilimnion and hypolimnion continues to decrease, winds are able to mix the vertical profile to greater depths. This process continues until the temperature is uniform throughout the basin (see Figure 3.9). Now, pond and lake water circulate throughout the basin. This process of vertical circulation, called the turnover, is an important component of nutrient dynamics in open-water ecosystems (see Chapter 21). Stirred by wind, the process of vertical mixing may last until ice forms at the surface. Then comes winter, and as the surface water cools to below 4°C, it becomes lighter again and remains on the surface. (Remember, water becomes lighter above and below 4°C; see Figure 3.4.) If the winter is cold enough, surface water freezes; otherwise, it remains close to 0°C. Now the warmest place in the pond or lake is on the bottom. In spring, the breakup of ice and heating of surface water with increasing inputs of solar radiation to the surface again causes the water to stratify. Because not all bodies of water experience such seasonal changes in stratification, this phenomenon is not necessarily characteristic of all deep bodies of water. In some deep lakes and the oceans, the thermocline simply descends during periods of turnover and does not disappear at all. In such bodies of water, the bottom water never becomes mixed with the top layer. In shallow lakes and ponds, temporary stratification of short duration may occur; in other bodies of water, stratification may exist, but the depth is not sufficient to develop a distinct thermocline. However, some form of thermal stratification occurs in all open bodies of water. Temperature and density profiles with water depth for an open body of water such as a lake or pond. (a) The vertical profile of temperature might be expected to resemble the profile of light presented in Figure 3.7, but vertical mixing of the surface waters transports heat to the waters below. Below this mixed layer, temperatures decline rapidly in a region called the thermocline. Below the thermocline, temperatures continue declining at a slower rate. The vertical profile can therefore be divided into three distinct zones: epilimnion, thermocline, and hypolimnion. (b) The rapid decline in temperature in the thermocline results in a distinct difference in water density (see Figure 3.4) in the warmer epilimnion as compared to the cooler waters of the hypolimnion, leading to a two-layer density profile—warm, low-density surface water and cold, high-density deep water. The temperature of a flowing body of water (stream or river), on the other hand, is variable (Figure 3.10). Small, shallow streams tend to follow, but lag behind, air temperatures. They warm and cool with the seasons but rarely fall below freezing in winter. Streams with large areas exposed to sunlight are warmer than those shaded by trees, shrubs, and high banks. That fact is ecologically important because temperature affects the stream community, influencing the presence or absence of cool- and warm-water organisms. For example, the dominant predatory fish shift from species such as trout and smallmouth bass, which require cooler water and more oxygen, to species such as suckers and catfish, which require warmer water and less oxygen (see Figure 24.13). 3.5 Water Functions as a Solvent As you stir a spoonful of sugar into a glass of water, it dissolves, forming a homogeneous, or uniform, mixture. A liquid that is a homogeneous mixture of two or more substances is called a solution. The dissolving agent of a solution is the solvent, and the substance that is dissolved is referred to as the solute. A solution in which water is the solvent is called an aqueous solution. Water is an excellent solvent that can dissolve more substances than can any other liquid. This extraordinary ability makes water a biologically crucial substance. Water provides a fluid that dissolves and transports molecules of nutrients and waste products, helps to regulate temperature, and preserves chemical equilibrium within living cells. The solvent ability of water is largely a result of the bonding discussed in Section 3.2. Because the H atom is bonded to the O atoms asymmetrically (see Figure 3.3), one side of every water molecule has a permanent positive charge and the other side has a permanent negative charge; such a situation is called a permanent dipole (where dipole refers to oppositely charged poles). Because opposite charges attract, water molecules are strongly attracted to one another; they also attract other molecules carrying a charge. Compounds that consist of electrically charged atoms or groups of atoms are called ions. Sodium chloride (table salt), for example, is composed of positively charged sodium ions (Na+) and negatively charged chloride ions (Cl–) arranged in a crystal lattice. When placed in water, the attractions between negative (oxygen atom) and positive (hydrogen atoms) charges on the water molecule (see Figure 3.3) and those of the sodium and chloride atoms are greater than the forces (ionic bonds) holding the salt crystals together. Consequently, the salt crystals readily dissociate into their component ions when placed in contact with water; that is, they dissolve. The solvent properties of water are responsible for the presence of most of the minerals (elements and inorganic compounds) found in aquatic environments. When water condenses to form clouds, it is nearly pure except for some dissolved atmospheric gases. In falling to the surface as precipitation, water acquires additional substances from particulates and dust particles suspended in the atmosphere. Water that falls on land flows over the surface and percolates into the soil, obtaining more solutes. Surface waters, such as streams and rivers, pick up more solvents from the substances through and over which they flow. The waters of most rivers and lakes contain 0.01–0.02 percent dissolved minerals. The relative concentrations of minerals in these waters reflect the substrates over which the waters flow. For example, waters that flow through areas where the underlying rocks consist largely of limestone, composed primarily of calcium carbonate (CaCO3), will have high concentrations of calcium (Ca2+) and bicarbonate (HCO3–). In contrast to freshwaters, the oceans have a much higher concentration of solutes. In effect, the oceans function as a large still. The flow of freshwaters into the oceans continuously adds to the solute content of the waters, as pure water evaporates from the surface to the atmosphere. The concentration of solutes, however, cannot continue to increase indefinitely. When the concentration of specific elements reaches the limit set by the maximum solubility of the compounds they form (grams per liter), the excess amounts will precipitate and be deposited as sediments. Calcium, for example, readily forms calcium carbonate (CaCO3) in the waters of the oceans. The maximum solubility of calcium carbonate, however, is only 0.014 gram per liter of water, a concentration that was reached early in the history of the oceans. As a result, calcium ions continuously precipitate out of solution and are deposited as limestone sediments on the ocean bottom. In contrast, the solubility of sodium chloride is high (360 grams per liter). In fact, these two elements, sodium and chlorine, make up some 86 percent of sea salt. Sodium and chlorine—along with other major elements such as sulfur, magnesium, potassium, and calcium, whose relative proportions vary little—compose 99 percent of sea salts (Figure 3.11). Determination of the most abundant element, chlorine, is used as an index of salinity. Salinity is expressed in practical salinity units (psu), represented as ‰ and measured as grams of chlorine per kilogram of water. The salinity of the open sea is fairly constant, averaging about 35‰. In contrast, the salinity of freshwater ranges from 0.065 to 0.30‰. However, over geologic timescales (hundreds of millions of years), the salinity of the oceans has increased and continues to do so. 3.6 Oxygen Diffuses from the Atmosphere to the Surface Waters Water’s role as a solvent is not limited to dissolving solids. The surface of a body of water defines a boundary with the atmosphere. Across this boundary, gases are exchanged through the process of diffusion. Diffusion is the general tendency of molecules to move from a region of high concentration to one of lower concentration. The process of diffusion results in a net transfer of two metabolically important gases, oxygen and carbon dioxide, from the atmosphere (higher concentration) into the surface waters (lower concentration) of aquatic environments Oxygen diffuses from the atmosphere into the surface water. The rate of diffusion is controlled by the solubility of oxygen in water and the steepness of the diffusion gradient (the difference in concentration between the air and the surface waters where diffusion occurs). The solubility of gases in water is a function of temperature, pressure, and salinity. The saturation value of oxygen is greater for cold water than warm water because the solubility (ability to stay in solution) of a gas in water decreases as the temperature rises. However, solubility increases as atmospheric pressure increases and decreases as salinity increases, which is not significant in freshwater. Once oxygen enters the surface water, the process of diffusion continues, and oxygen diffuses from the surface to the waters below (because of their lower concentration). Water, with its greater density and viscosity relative to air, limits how quickly gases diffuse through it. Gases diffuse some 10000 times slower in water than in air. In addition to the process of diffusion, oxygen absorbed by surface water is mixed with deeper water by turbulence and internal currents. In shallow, rapidly flowing water and in wind-driven sprays, oxygen may reach and maintain saturation and even supersaturated levels because of the increase of absorptive surfaces at the air–water interface. Oxygen is lost from the water as temperatures rise, decreasing solubility, and through the uptake of oxygen by aquatic life. Oxygen stratification in Mirror Lake, New Hampshire, in winter, summer, and late fall. The late fall turnover results in uniform temperature as well as uniform distribution of oxygen throughout the lake basin. In summer, a pronounced stratification of both temperature and oxygen exists. Oxygen declines sharply in the thermocline and is nonexistent on the bottom because of its uptake by decomposer organisms in the sediments. In winter, oxygen levels are high in surface water reflecting higher solubility. Formation of ice during winter, however, can greatly reduce diffusion into surface waters. (Adapted from Likens 1985.) During the summer, oxygen, like temperature (see Section 3.4), may become stratified in lakes and ponds. The amount of oxygen is usually greatest near the surface, where an interchange between water and atmosphere, further stimulated by the stirring action of the wind, takes place (Figure 3.12). Besides entering the water by diffusion from the atmosphere, oxygen is also a product of photosynthesis, which is largely restricted to the surface waters because of the limitations of available light (see Figure 3.7 and Chapter 6). The quantity of oxygen decreases with depth because of the oxygen demand of decomposer organisms living in the bottom sediments (Chapter 21). During spring and fall turnover, when water recirculates through the lake, oxygen becomes replenished in deep water. In winter, the reduction of oxygen in unfrozen water is slight because the demand for oxygen by organisms is reduced by the cold, and oxygen is more soluble at low temperatures. Under ice, however, oxygen depletion may be serious as a result of the lack of diffusion from the atmosphere to the surface waters. As with ponds and lakes, oxygen is not distributed uniformly within the depths of the oceans (Figure 3.13). A typical oceanic oxygen profile shows a maximum amount in the upper 10–20 m, where photosynthetic activity and diffusion from the atmosphere often lead to saturation. With increasing depth, oxygen content declines. In the open waters of the ocean, concentrations reach a minimum value of 500–1000 m, a region referred to as the oxygen minimum zone. Unlike lakes and ponds, where the seasonal breakdown of the thermocline and resultant mixing of surface and deep waters result in a dynamic gradient of temperature and oxygen content, the limited depth of surface mixing in the deep oceans maintains the vertical gradient of oxygen availability year-round. Vertical profile of oxygen with depth in the Atlantic Ocean. The oxygen content of the waters declines to a depth known as the oxygen minimum zone. Oxygen increase below this may be the result of the influx of cold, oxygen-rich waters that sank in the polar waters. The availability of oxygen in aquatic environments characterized by flowing water is different. The constant churning of stream water over riffles and falls gives greater contact with the atmosphere; the oxygen content of the water is high, often near saturation for the prevailing temperature. Only in deep holes or in polluted waters does dissolved oxygen show any significant decline (see Chapter 24, Ecological Issues & Applications). Even under ideal conditions, gases are not very soluble in water. Rarely is oxygen limited in terrestrial environments. In aquatic environments, the supply of oxygen, even at saturation levels, is meager and problematic. Compared with its concentration of 0.21 liter per liter in the atmosphere (21 percent by volume), oxygen in water reaches a maximum solubility of 0.01 liter per liter (1 percent) in freshwater at a temperature of 0°C. As a result, the concentration of oxygen in aquatic environments often limits respiration and metabolic activity. 3.7 Acidity Has a Widespread Influence on Aquatic Environments The solubility of carbon dioxide is somewhat different from that of oxygen in its chemical reaction with water. Water has a considerable capacity to absorb carbon dioxide, which is abundant in both freshwater and saltwater. Upon diffusing into the surface, carbon dioxide reacts with water to produce carbonic acid (H2CO3): CO2+H2O⇋H2CO3CO2 + H2O⇋ H2CO3 Carbonic acid further dissociates into a hydrogen ion and a bicarbonate ion: H2CO3⇋HCO3−+H+H2CO3 ⇋ HCO3− +H+ Bicarbonate may further dissociate into another hydrogen ion and a carbonate ion: HCO−3⇋H++CO2−3HCO3− ⇋ H++CO32− The carbon dioxide–carbonic acid–bicarbonate system is a complex chemical system that tends to stay in equilibrium. (Note that the arrows in the preceding equations go in both directions.) Therefore, if carbon dioxide (CO2) is removed from the water, the equilibrium is disturbed and the equations will shift to the left, with carbonic acid and bicarbonate producing more CO2 until a new equilibrium is reached. The chemical reactions just described result in the production and absorption of free hydrogen ions (H+). The abundance of hydrogen ions in solution is a measure of acidity. The greater the number of H+ ions, the more acidic is the solution. Alkaline solutions are those that have a large number of OH– (hydroxyl ions) and few H+ ions. The measurement of acidity and alkalinity is called pH, calculated as the negative logarithm (base 10) of the concentration of hydrogen ions in solution. In pure water, a small fraction of molecules dissociates into ions: H2O → H+ + OH–, and the ratio of H+ ions to OH– ions is 1:1. Because both occur in a concentration of 10–7 moles per liter, a neutral solution has a pH of 7 [–log(10–7) = 7]. A solution departs from neutral when the concentration of one ion increases and the other decreases. Customarily, we use the negative logarithm of the hydrogen ion to describe a solution as an acid or a base. Thus, a gain of hydrogen ions to 10–6 moles per liter means a decrease of OH– ions to 10–8 moles per liter, and the pH of the solution is 6. The negative logarithmic pH scale goes from 1 to 14. A pH greater than 7 denotes an alkaline solution (greater OH– concentration) and a pH of less than 7 an acidic solution (greater H+ concentration). Although pure water is neutral in pH, because the dissociation of the water molecule produces equal numbers of H+ and OH– ions, the presence of CO2 in the water alters this relationship. The preceding chemical reactions result in the production and absorption of H+ ions. Because the abundance of hydrogen ions in solution is the measure of acidity, the dynamics of the carbon dioxide–carbonic acid–bicarbonate system directly affect the pH of aquatic ecosystems. In general, the carbon dioxide–carbonic acid–bicarbonate system functions as a buffer to keep the pH of water within a narrow range. It does this by absorbing hydrogen ions in the water when they are in excess (producing carbonic acid and bicarbonates) and producing them when they are in short supply (producing carbonate and bicarbonate ions). At neutrality (pH 7), most of the CO2 is present as HCO3– ( Figure 3.14). At a high pH, more CO2 is present as CO32– than at a low pH, where more CO2 occurs in the free condition. Addition or removal of CO2 affects pH, and a change in pH affects CO2. The pH of natural waters ranges between 2 and 12. Waters draining from watersheds dominated geologically by limestone (CaCO3) have a much higher pH and are well buffered as compared to waters from watersheds dominated by acid sandstone and granite. The presence of the strongly alkaline ions sodium, potassium, and calcium in ocean waters results in seawater being slightly alkaline, usually ranging from 7.5 to 8.4. The pH of aquatic environments can exert a powerful influence on the distribution and abundance of organisms. Increased acidity can affect organisms directly, by influencing physiological processes, and indirectly, by influencing the concentrations of toxic heavy metals. Tolerance limits for pH vary among plant and animal species, but most organisms cannot survive and reproduce at a pH below about 4.5. Aquatic organisms are unable to tolerate low pH conditions largely because acidic waters contain high concentrations of aluminum. Aluminum is highly toxic to many species of aquatic life and thus leads to a general decline in aquatic populations. Theoretical percentages of carbon dioxide (CO2) in each of the three forms present in water in relation to pH. At low values of pH (acidic conditions), most of the CO2 is in its free form. At intermediate values (neutral conditions) bicarbonate dominates, whereas under alkaline conditions most of the CO2 is in the form of carbonate ions. Interpreting Ecological Data Q1. Under conditions of neutral pH, what is the relative abundance of the different forms of CO2? Q2. The current pH of the ocean is approximately 8.1. what is the dominant form of CO2? Aluminum is insoluble when the pH is neutral or basic. Insoluble aluminum is present in high concentrations in rocks, soils, and river and lake sediments. Under normal pH conditions, the aluminum concentrations of lake water are low; however, as the pH drops and becomes more acidic, aluminum begins to dissolve, raising the concentration in solution. 3.8 Water Movements Shape Freshwater and Marine Environments The movement of water—currents in streams and waves in an open body of water or breaking on a shore—determines the nature of many aquatic environments. The velocity of a current molds the character and structure of a stream. The shape and steepness of the stream channel, its width, depth, and roughness of the bottom, and the intensity of rainfall and rapidity of snowmelt all affect velocity. In fast streams, velocity of flow is 50 cm per second or higher (see Chapter 24, Quantifying Ecology 24.1). At this velocity, the current removes all particles less than 5 millimeters (mm) in diameter and leaves behind a stony bottom. High water volume increases the velocity; it moves bottom stones and rubble, scours the streambed, and cuts new banks and channels. As the gradient decreases and the width, depth, and volume of water increase, silt and decaying organic matter accumulate on the bottom. Thus, the stream’s character changes from fast water to slow (Figure 3.15). Wind generates waves on large lakes and on the open sea. The frictional drag of the wind on the surface of smooth water causes ripples. As the wind continues to blow, it applies more pressure to the steep side of the ripple, and wave size begins to grow. As the wind becomes stronger, short, choppy waves of all sizes appear; as they absorb more energy, they continue to grow. When the waves reach a point where the energy supplied by the wind equals the energy lost by the breaking waves, they become whitecaps. Up to a certain point, the stronger the wind, the higher the waves. The waves breaking on a beach do not contain water driven in from distant seas. Each particle of water remains largely in the same place and follows an elliptical orbit with the passage of the wave. As a wave moves forward, it loses energy to the waves behind and disappears, its place taken by another. The swells breaking on a beach are distant descendants of waves generated far out at sea. As the waves approach land, they advance into increasingly shallow water. When the bottom of the wave intercepts the ocean floor, the wavelength shortens and the wave steepens until it finally collapses forward, or breaks. As the waves break onshore, they dissipate their energy, pounding rocky shores or tearing away sandy beaches in one location and building up new beaches elsewhere. We have discussed the patterns of ocean currents, influenced by the direction of the prevailing winds and the Coriolis effect in Chapter 2 (see Section 2.4). As the warm surface currents of the tropical waters move northward and southward (see map of surface currents in Figure 2.13), they bring up deep, cold, oxygenated waters from below, a process known as upwelling (Figure 3.16a). A similar pattern occurs in coastal regions. Winds blowing parallel to the coast move the surface waters offshore. Water moving upward from the deep replaces this surface water, creating a pattern of coastal upwelling (Figure 3.16b). Along the equator, the Coriolis effect acts to pull the westward-flowing currents to the north and south (purple solid arrows), resulting in an upwelling of deeper cold waters to the surface. (b) Along the western margins of the continents, the Coriolis effect causes the surface waters to move offshore (purple solid arrows). Movement of the surface waters offshore results in an upwelling of deeper, colder waters to the surface. Example shown is for the Northern Hemisphere. 3.9 Tides Dominate the Marine Coastal Environment Tides profoundly influence the rhythm of life on ocean shores. Tides result from the gravitational pulls of the Sun and the Moon, each of which causes two bulges (tides) in the waters of the oceans. The two bulges caused by the Moon occur at the same time on opposite sides of Earth on an imaginary line extending from the Moon through the center of Earth (Figure 3.17). The tidal bulge on the Moon side is a result of gravitational attraction; the bulge on the opposite side occurs because the gravitational force there is less than at the Earth’s center. As Earth rotates eastward on its axis, the tides advance westward. Thus, in the course of one daily rotation, Earth passes through two of the lunar tidal bulges, or high tides, and two of the lows, or low tides, at right angles (90° longitude difference) to the high tides. Tides result from the gravitational pull of the Moon. Centrifugal force applied to a kilogram of mass is 3.38 milligrams (mg). This force on a rotating Earth is balanced by gravitational force, except at moving points on Earth’s surface that are directly aligned with the Moon. Thus, the centrifugal force at point N, the center of the rotating Earth, is 3.38 mg. Point T is directly aligned with the Moon. At this point, the Moon’s gravitational force is 3.49 mg, a difference of 0.11 mg. Because the Moon’s gravitational force is greater than the centrifugal force at T, the force is directed away from the Earth and causes a tidal bulge. At point A, the Moon’s gravitational force is 3.27 mg, 0.11 mg less than the centrifugal force at N. This causes a tidal bulge on the opposite side of Earth. The Sun also causes two tides on opposite sides of Earth, and these tides have a relation to the Sun like that of the lunar tides to the Moon. Because the Sun has a weaker gravitational pull than the Moon does, solar tides are partially masked by lunar tides—except for two times during the month: when the Moon is full and when it is new. At these times, Earth, Moon, and Sun are nearly in line, and the gravitational pulls of the Sun and the Moon are additive. This combination makes the high tides of those periods exceptionally large, with maximum rise and fall. These are the fortnightly spring tides, a name derived from the Saxon word sprungen, which refers to the brimming fullness and active movement of the water. When the Moon is at either quarter, its pull is at right angles to the pull of the Sun, and the two forces interfere with each other. At those times, the differences between high and low tides are exceptionally small. These are called the neap tides, from an old Scandinavian word meaning “barely enough.” Tides are not entirely regular, nor are they the same all over Earth. They vary from day to day in the same place, following the waxing and waning of the Moon. They may act differently in several localities within the same general area. In the Atlantic, semidaily tides are the rule. In the Gulf of Mexico and the Aleutian Islands of Alaska, the alternate highs and lows more or less cancel each other out, and flood and ebb follow one another at about 24-hour intervals to produce one daily tide. Mixed tides in which successive or low tides are of significantly different heights through the cycle are common in the Pacific and Indian oceans. These tides are combinations of daily and semidaily tides in which one partially cancels out the other. Local tides around the world are inconsistent for many reasons. These reasons include variations in the gravitational pull of the Moon and the Sun as a result of the elliptical orbit of Earth, the angle of the Moon in relation to the axis of Earth, onshore and offshore winds, the depth of water, the contour of the shore, and wave action. The area lying between the water lines of high and low tide, referred to as the intertidal zone, is an environment of extremes. The intertidal zone undergoes dramatic shifts in environmental conditions with the daily patterns of inundation and exposure. As the tide recedes, the uppermost layers of life are exposed to air, wide temperature fluctuations, intense solar radiation, and desiccation for a considerable period, whereas the lowest fringes of the tidal zone may be exposed only briefly before the high tide submerges them again. Temperatures on tidal flats may rise to 38°C when exposed to direct sunlight and drop to 10°C within a few hours when the flats are covered by water. Organisms living in the sand and mud do not experience the same violent temperature fluctuations as those living on rocky shores do. Although the surface temperature of the sand at midday may be 10°C (or more) higher than that of the returning seawater, the temperature a few centimeters below the sand’s surface remains almost constant throughout the year (see Section 25.3). 3.10 The Transition Zone between Freshwater and Saltwater Environments Presents Unique Constraints Water from streams and rivers eventually drains into the sea. The place where freshwater mixes with saltwater is called an estuary. Temperatures in estuaries fluctuate considerably, both daily and seasonally. Sun and inflowing and tidal currents heat the water. High tide on the mudflats may heat or cool the water, depending on the season. The upper layer of estuarine water may be cooler in winter and warmer in summer than the bottom—a condition that, as in a lake, will cause spring and autumn turnovers (see Figures 3.9 and 3.12). In the estuary, where freshwater meets the sea, the interaction of inflowing freshwater and tidal saltwater influences the salinity of the estuarine environment. Salinity varies vertically and horizontally, often within one tidal cycle (Figure 3.18). Vertical and horizontal stratification of salinity from the river mouth to the estuary at high tide (brown lines) and low tide (blue lines). At high tide, the incoming seawater increases the salinity toward the river mouth. At low tide, salinity is reduced. Salinity increases with depth because lighter freshwater flows over denser saltwater. Salinity may be the same from top to bottom or it may be completely stratified, with a layer of freshwater on top and a layer of dense, salty water on the bottom. Salinity is homogeneous when currents are strong enough to mix the water from top to bottom. The salinity in some estuaries is homogeneous at low tide, but at high tide a surface wedge of seawater moves upstream more rapidly than the bottom water. Salinity is then unstable, and density is inverted. The seawater on the surface tends to sink as lighter freshwater rises, and mixing takes place from the surface to the bottom. This phenomenon is known as tidal overmixing. Strong winds, too, tend to mix saltwater with freshwater in some estuaries, but when the winds are still, the river water flows seaward on a shallow surface over an upstream movement of seawater, more gradually mixing with the salt. Horizontally, the least saline waters are at the river mouth and the most saline at the sea (see Figure 3.18). Incoming and outgoing currents deflect this configuration. In all estuaries of the Northern Hemisphere, outward-flowing freshwater and inward-flowing seawater are deflected to the right (relative to the axis of water flow from the river to ocean) because of Earth’s rotation (Coriolis effect; see Section 2.3). As a result, salinity is higher on the left side; the concentration of metallic ions carried by rivers varies from drainage to drainage; and salinity and chemistry differ among estuaries. The portion of dissolved salts in the estuarine waters remains about the same as that of seawater, but the concentration varies in a gradient from freshwater to sea. To survive in estuaries, aquatic organisms must have evolved physiological or behavioral adaptations to changes in salinity. Many oceanic species of fish are able to move inward during periods when the flow of freshwater from rivers is low and the salinity of estuaries increases. Conversely, freshwater fish move into the estuarine environment during periods of flood when salinity levels drop. Because of the stressful conditions that organisms face in the mixed zones of estuaries, there is often a relatively low diversity of organisms despite the high productivity found in these environments (see Chapter 24). Ecological Issues & Applications Rising Atmospheric Concentrations of CO2 Are Impacting Ocean Acidity The exchange of carbon dioxide (CO2) between the atmosphere and the surface waters of the oceans is governed by the process of diffusion, with the net exchange moving CO2 from higher concentrations (atmosphere) to lower concentrations (surface waters) (Section 3.7). Upon diffusing into the surface, the CO2 reacts with the water to produce carbonic acid (H2CO3), which further dissociates into a hydrogen ion (H+) and a bicarbonate ion (HCO3−). The bicarbonate may further dissociate into another hydrogen ion and a carbonate ion (CO32−). In both of these chemical reactions, free hydrogen ions (H+) are produced, the abundance of which is a measure of acidity. The greater the number of H+ ions, the lower the value of pH and the more acidic the solution. Under current ocean conditions, about 89 percent of the carbon dioxide dissolved in seawater takes the form of a bicarbonate ion, about 10 percent as a carbonate ion, and 1 percent as dissolved gas, and the pH of seawater on the surface of the oceans has remained relatively steady for millions of years at a value of about 8.2 (slightly basic—7.0 is neutral; see Figure 3.14). Since the height of the Industrial Revolution in the 19th century, however, atmospheric concentrations of CO2 have been steadily rising as a result of the burning of fossil fuels (see Chapter 2, Ecological Issues & Applications and Chapter 27). As a consequence, the diffusion gradient of CO2 between the atmosphere and oceans has increased, resulting in an increasing uptake of CO2 into the surface waters. As a consequence, the pH of the surface waters of the oceans has fallen by about 0.1 pH unit from preindustrial times to today (Figure 3.19). Recall from Section 3.7 that the pH scale is logarithmic (log10; thus for every drop of 1 pH unit, hydrogen ion levels increase by a factor of 10), so this 0.1-unit drop in pH is equivalent to about a 25 percent increase in the ocean hydrogen ion concentration. According to estimates from the Intergovernmental Panel on Climate Change (IPCC; see Chapter 2, Ecological Issues & Applications), under the expected trajectory of fossil fuel use and rising atmospheric CO2 concentrations, pH is likely to drop by 0.3–0.4 units by the end of the 21st century and increase ocean hydrogen ion concentration (or acidity) by 100 to 150 percent. Increased absorption of CO2 by the surface waters of the oceans can potentially impact life in the oceans in a variety of ways, both positive and negative. Photosynthetic algae and plants may benefit from higher CO2 concentrations in the surface waters because elevated CO2 may enhance rates of photosynthesis (see Chapter 6, Ecological Issues & Applications). On the other hand, one of the most important negative impacts of increasing ocean acidity relates to the process of calcification—the production of shells and plates out of calcium carbonate (CaCO3)—which is important to the biology and survival of a wide range of marine organisms. CaCO3 is formed in marine environments through the reaction of calcium and carbonate ions: CO32−+C2+a⇋CaCO3CO32− + C2+a⇋ CaCO3 As with the chemical equations describing the formation of bicarbonate and carbonate ions from dissolved CO2, the reaction involved in the formation of CaCO3 proceeds in both directions. As sea water pH declines (acidity increases), carbonate ions (CO32−) function like an antacid to neutralize the H+, forming more bicarbonate (CO32− + H+ ↔ HCO3−). Therefore, declining pH results in an associated decline in carbonate ion concentrations (Figure 3.19a). This decline in carbonate ion concentration shifts the preceding equation in favor of the disassociation of CaCO3 minerals into calcium and carbonate ions. The resulting decline in dissolved CaCO3 minerals can have a significant impact on calcifying species, including oysters, clams, sea urchins, shallow water corals, deep sea corals, and calcareous plankton. The process of calcification by marine organisms involves the precipitation of dissolved CaCO3 into solid CaCO3 structures, such as coccoliths (individual plates of CaCO3 formed by single-celled algae; Figure 3.20). After they are formed, these structures are vulnerable to dissolution unless the surrounding seawater contains saturating concentrations of CaCO3. As carbonate ions become depleted because of declining pH, seawater becomes undersaturated with respect to two CaCO3 minerals vital for calcification, aragonite and calcite (Figure 3.19b and 3.19c). Current estimates suggest that the oceans are becoming undersaturated with respect to aragonite at the poles, where the cold and dense waters most readily absorb atmospheric CO2, and that under projected rates of CO2 emissions (IPCC), undersaturation would extend throughout the entire Southern Ocean (<60° S) and into the subarctic Pacific by the end of the century (2100). These changes will threaten high-latitude aragonite secreting organisms including cold-water corals, which provide essential fish habitat, and shelled pteropods (free-swimming pelagic sea snails and sea slugs; Figure 3.21), an abundant food source for marine predators In a review of experimental studies that have examined the response of marine calcifying species to elevated CO2 conditions, Scott Doney of Woods Hole Oceanographic Institute and colleagues found that the degree of sensitivity varies among species, and some species may even show enhanced calcification at elevated CO2 levels (Table 3.1). The researchers found, however, that in the vast majority of species, including every study published that has examined the calcification rates of coral species, elevated CO2 concentrations and the associated decreasing aragonite saturation state had a negative effect on calcification rates. The researchers concluded that “ocean acidification impacts processes so fundamental to the overall structure and function of marine ecosystems that any significant changes could have far-reaching consequences for the oceans of the future.” Summary The Water Cycle 3.1 Water follows a cycle, traveling from the air to Earth and returning to the atmosphere. It moves through cloud formation in the atmosphere, precipitation, interception, and infiltration into the ground. It eventually reaches groundwater, springs, streams, and lakes from which evaporation takes place, bringing water back to the atmosphere in the form of clouds. The various aquatic environments are linked, either directly or indirectly, by the water cycle. The largest reservoir in the global water cycle is the oceans, which contain more than 97 percent of the total volume of water on Earth. In contrast, the atmosphere is one of the smallest reservoirs but has a fast turnover time. Properties of Water 3.2 Water has a unique molecular structure. The hydrogen atoms are located on the side of the water molecule that has a positive charge. The opposite side, where the oxygen atom is located, has a negative charge, thus polarizing the water molecule. Because of their polarity, water molecules become coupled with neighboring water molecules to produce a lattice-like structure with unique properties. Depending on its temperature, water may occur in the form of a liquid, solid, or gas. It absorbs or releases considerable quantities of heat with a small rise or fall in temperature. Water has a high viscosity that affects its flow. It exhibits high surface tension, caused by a stronger attraction of water molecules for each other than for the air above the surface. If a body is submerged in water and weighs less than the water it displaces, it is subjected to the upward force of buoyancy. These properties are important ecologically and biologically. Light 3.3 Both the quantity and quality of light change with water depth. In pure water, red and infrared light are absorbed first, followed by yellow, green, and violet; blue penetrates the deepest. Temperature in Aquatic Environments 3.4 Lakes and ponds experience seasonal shifts in temperature. In summer there is a distinct vertical gradient of temperature, resulting in a physical separation of warm surface waters and the colder waters below the thermocline. When the surface waters cool in the fall, the temperature becomes uniform throughout the basin and water circulates throughout the lake. A similar mixing takes place in the spring when the water warms. In some deep lakes and the oceans, the thermocline simply descends during turnover periods and does not disappear at all. Temperature of flowing water is variable, warming and cooling with the season. Within the stream or river, temperatures vary with depth, amount of shading, and exposure to sun. Water as a Solvent 3.5 Water is an excellent solvent with the ability to dissolve more substances than any other liquid can. The solvent properties of water are responsible for most of the minerals found in aquatic environments. The waters of most rivers and lakes contain a relatively low concentration of dissolved minerals, determined largely by the underlying bedrock over which the water flows. In contrast, the oceans have a much higher concentration of solutes. As pure water evaporates from the surface to the atmosphere, the flow of freshwaters into the oceans continuously adds to the solute content of the waters. The solubility of sodium chloride is high; together with chlorine, it makes up some 86 percent of sea salt. The concentration of chlorine is used as an index of salinity. Salinity is expressed in practical salinity units (psu; represented as ‰, measured as grams of chlorine per kilogram of water). Oxygen 3.6 Oxygen enters the surface waters from the atmosphere through the process of diffusion. The amount of oxygen water can hold depends on its temperature, pressure, and salinity. In lakes, oxygen absorbed by surface water mixes with deeper water by turbulence. During the summer, oxygen may become stratified, decreasing with depth because of decomposition in bottom sediments. During spring and fall turnover, oxygen becomes replenished in deep water. Constant swirling of stream water gives it greater contact with the atmosphere and thus allows it to maintain a high oxygen content. Acidity 3.7 The measurement of acidity is pH, the negative logarithm of the concentration of hydrogen ions in solution. In aquatic environments, a close relationship exists between the diffusion of carbon dioxide into the surface waters and the degree of acidity and alkalinity. Acidity influences the availability of nutrients and restricts the environment of organisms sensitive to acid situations. Water Movement 3.8 Currents in streams and rivers as well as waves in open sea and breaking on ocean shores determine the nature of many aquatic and marine environments. The velocity of currents shapes the environment of flowing water. Waves pound rocky shores and tear away and build up sandy beaches. Movement of water in surface currents of the ocean affects the patterns of deep-water circulation. As the equatorial currents move northward and southward, deep waters move up to the surface, forming regions of upwelling. In coastal regions, winds blowing parallel to the coast create a pattern of coastal upwelling. Tides 3.9 Rising and falling tides shape the environment and influence the rhythm of life in coastal intertidal zones. Estuaries 3.10 Water from all streams and rivers eventually drains into the sea. The place where this freshwater joins and mixes with the salt is called an estuary. Temperatures in estuaries fluctuate considerably, both daily and seasonally. The interaction of inflowing freshwater and tidal saltwater influences the salinity of the estuarine environment. Salinity varies vertically and horizontally, often within one tidal cycle. Ocean Acidification Ecological Issues & Applications Rising atmospheric concentrations of carbon dioxide have resulted in increased concentrations in the surface waters of the oceans. The increased carbon dioxide concentrations of the surface waters have resulted in a decline in pH and reduced carbonate concentrations. The reduction in carbonate concentrations has reduced calcium carbonate mineral concentrations that are essential for calcifying marine species. How to Measure and Use Leaf Area Index 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.
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