Keywords: carbon, nitrogen, phosphorus, element cycles, photosynthesis, respiration, decomposition, biogeochemistry, net primary production, net ecosystem production, contaminants, reduction-oxidation, rock weathering, global warming, methane, nitrous oxide

1. Introduction to the Element Cycles

2. Brief History of the Elements

3. Mass Flow of Air, Water, and Rock

4. The Carbon Cycle

5. Nutrients and Limiting Elements

6. Cycling of Radiatively Active Gases

7. Human Influence on Global Biogeochemical Cycles

Related Chapters

Glossary

Bibliography

Biographical Sketches

Life is manufactured from a small number of elements arranged to form an immense variety of compounds. Organisms require elements in fairly constant ratios, and the rate of supply limits organic productivity. Because the ultimate sources of life-supporting elements are Earths crust, water, and the atmosphere, biological processes are tightly linked to the exchange of elements between these reservoirs, some of which cycle slowly. Organic evolution has improved element supply rates through novel physiological adaptations that enhance element acquisition and nutrient use. Some scientists argue that evolution has favored life forms with characteristics that stabilize Earths climate system, but this hypothesis may not apply to Homo sapiens who are presently changing climate by burning fossil fuels, deforestation and a variety of other activities.

The carbon cycle is an ideal example of how biological, chemical, and physical processes interact to influence the capacity of Earth to support life. Carbon in the atmosphere exists mostly as CO₂ and CH₄, both of which are greenhouse gases. Plants and phytoplankton consume about 30% of the CO₂ in the atmosphere annually. Much of the carbon removed from the atmosphere is returned quickly through respiration, but a portion remains stored for decades to millennia in plant biomass, soil organic matter and marine sediments. Exchange between these reservoirs and the atmosphere can explain many characteristics of the climate system. Human activity has greatly enhanced such exchanges.

Nitrogen and phosphorus are the nutrient elements that most often limit organic productivity. The atmosphere and rocks are the ultimate (long-term) sources of nutrients. Yet, most of the annual demand for nutrients is met through efficient recycling mechanisms. Life itself exercises strong control over nutrient element cycling.

Many trace elements are both nutrients and contaminants depending on their concentration. At low concentrations they may limit productivity due to their absence, while at high concentrations they suppress productivity because they are toxic. Their ecological importance is generally greater than their abundance would suggest.

In December 1990, the Galileo spacecraft observed a planet in our solar system with oceans of liquid water, a ubiquitous green pigment, and an atmosphere with O₂ and methane in extreme thermodynamic disequilibrium. Of course, the planet was Earth. This experiment was conceived by Carl Sagan as an opportunity to test a key characteristic of life at the global scale n its influence on the chemical forms and distribution of elements. Yet, life is just one force shaping global element cycles. Purely physical and chemical processes generate element cycles in the absence of life. Such processes include the input of solar radiation from the sun, mixing of the oceans and tectonic cycling of the crust (see The Geosphere). The study of natural element cycles has developed into the field of biogeochemistry, an area of scientific inquiry that seeks to integrate the traditional disciplines of biology, geology, and chemistry. A full appreciation of the field also requires knowledge of the physics of atmospheres, oceans, and continents.

Biogeochemistry involves the study of element sources, their transport, transformations, and sinks. Elements are constantly moving between reservoirs (e.g. soils, oceans, and atmospheres) and/or changing form (e.g., from carbon bound in CO₂ gas to carbon bound in sugar). Because most elements on Earth are neither created nor destroyed by processes operating on the planet, they pass through any given reservoir or chemical form repeatedly over time and are thereby cycled. Each process that modifies the form of an element may be considered in isolation, but in reality theses processes are coupled to an element cycle. Organisms participate in most phases of element cycles, but they are particularly strong element transformers (i.e., they change elements from one chemical form to another).

A key challenge facing biogeochemists is the need to integrate information across vast scales of time and space. Many element cycling studies are done at a scale of square meters, yet the results are most relevant when extrapolated to landscape or global scales measured in units ranging from 10 km² to 1 10⁶ km². Scales of time and space are often linked. For example, during photosynthesis light energy is rapidly converted to chemical energy in a process that requires a pH difference across membranes in the chloroplast (Figure 1). At the other end of the spectrum, yearly rates of forest growth depend upon the number and distribution of light harvesting leaves across large tracts of forest. One approach to scaling such information is to model the essential physical and biogeochemical processes with mathematical equations. Parameters in the equations are given values derived from small-scale studies or remotely sensed data. The equations are then used to predict chemical cycling at larger scales. Biogeochemical models benefit from spatially explicit databases of important attributes such as temperature, rainfall, precipitation, vegetation, and soil texture. They are often coupled to climate models to predict future climate change.

Figure 1. A framework for integrating scales of time and space ranging from the molecule to the biosphere for processes related to CO₂ uptake by vegetation. Source: Osmond C.B. and W.S. Chow (1988) Ecology of photosynthesis in the sun and shade: summary and prognostications. Australian Journal of Plant Physiology 15:19.

Models are also used to integrate across various scales of time ranging from years to millennia. The parameters for such models require information about past climates or distributions of plant communities. Fortunately, the history of chemical cycling on Earth is recorded in rocks, ice caps, coral reefs, tree rings, soils, aquatic sediments, and many other features. Sherlock Holmes, the fictitious English detective, would have been proud of the creative approaches that have been used to decipher and interpret these historical records.

This article briefly reviews the biological and physical forces that operate on element cycles. It focuses on carbon (C), nitrogen (N), and phosphorus (P) because of their key roles in energy acquisition and growth, and because they represent a range of possibilities for the importance of biological versus geological processes. Certain key features of the other element cycles are considered here, while more specific information is available in the articles in this section.

The pool of elements in which life exists today is largely the same as that endowed to Earth upon its formation about 4.6 Ga (see Universe as Earths Environment). However, there have been changes in the chemical forms and distribution of elements over time. The planet can be divided into three concentric spheres on the basis of differences in chemical composition: the core, mantle, and crust (outermost). Heavy elements, especially iron (Fe), are concentrated in the core, while the crust is enriched in relatively light elements that form aluminosilicate minerals such as potassium feldspar (KAlSi₃O₈). Thus, the biosphere arose in a chemical mixture that was unique to the planet's surface.

The atmosphere is composed of elements that were released as gases from molten rocks while the planet was still relatively hot. This process continues today with volcanic eruptions rich in H₂O, CO₂, SO₂, H₂S, and HCl. Volatile compounds that do not readily dissolve in oceans, such as N₂ and O₂, have accumulated to high levels in the atmosphere. These gases currently account for 78% and 21% of the atmosphere, respectively. Many ocean salts were derived from the dissolution of gases with high water solubility such as CO₂, SO₂, and Cl₂.

The oceans and atmosphere remained moderately reducing until microbes evolved O₂-generating photosynthesis as early as 3.5 Ga. Only much later did O₂ began to increase in the atmosphere. It was scrubbed from seawater in oxidation reactions with Fe²⁺ (a reduced, soluble form of Fe), and from the atmosphere by reduced gases such as CH₄. The oxidation of Fe²⁺ produces Fe(OH)₃ and related compounds that precipitate in water. This process probably accounts for the Banded Iron Formation, a geological feature that holds major economic deposits of iron ore in the United States. Peak deposition of the Banded Iron Formation occurred 2.53.0 Ga. The accumulation of O₂ in the atmosphere started a process of oxidizing Fe²⁺ and other reduced minerals in continental rocks.

As the supply of reduced minerals, solutes, and gases capable of reacting with O₂ dwindled, O₂ began to accumulate to high levels in the atmosphere; Earth's surface has been a highly oxidized environment ever since. The abundance of O₂ changed element cycling in several fundamental respects. For example, elements such as Fe and Mn typically precipitate in O₂-rich environments and therefore have low concentrations in water. Microorganisms that cannot tolerate O₂ now exist only in anaerobic environments such as wetland soils or the digestive tracts of termites and cattle. Organisms that use O₂ for respiration dominate the planet because the process yields far more energy than other forms of respiration. The abundance of O₂ allowed the formation of a protective ozone layer (O₃) that presumably enhanced the stability of organic chemicals, including the organic structure of the genetic code, by shielding organisms from high-energy ultraviolet radiation. Also, high O₂ levels made fire an important biogeochemical force. Because most of the O₂ in the atmosphere is balanced by reduced compounds buried in ocean basins, no amount of burning or forest destruction could significantly lower the atmospheres current O₂ content.

Element cycles are inextricably linked to the mass flow of the reservoirs in which they reside. The principle reservoirs are the atmosphere, ocean and crust. Each reservoir has a characteristic rate of change that influences rates of element cycling. Because all elements move between these reservoirs, their cycles may have both rapid and slow phases (see Mass and Energy).

The atmosphere is the most rapidly cycling reservoir with a mixing time of about three years. The portion of the atmosphere between the ground and a height of 1015 km is the troposphere, which holds 80% of the atmospheres mass. The stratosphere begins at the top of the troposphere and has a height of ~30 km. It has far less mass than the troposphere, but holds 90% of the atmosphere's ozone (O₃). Mixing between the two layers is generally slow. Gases and aerosols emitted at Earth's surface are likely to be removed before reaching the stratosphere due to chemical reactions and dissolution in rainfall. However, gases that are highly insoluble in water and chemically stable, such as CH₄, N₂O, and the chlorofluorocarbons (CFCs), are important exceptions to this rule. Because they are long-lived, these gases mix slowly into the stratosphere where they influence O₃ and climate (see Atmosphere and Climate).

The oceans hold >97% of the Earth's water and constitute large reservoirs of many key elements. As with the atmosphere, the oceans are composed of horizontal layers with characteristic rates and modes of element cycling. The mixed layer extends from the surface to between 75 and 200 m and is well mixed by winds. Gases, nutrients, organic matter, and heat that enter the ocean's surface will mix down to about 100 m depth in one year. Although the mixed-layer is <5% of the ocean's volume, it exchanges gases rapidly with the atmosphere and supports the productivity of phytoplankton. Below the surface lies the deep-water layer, which is 95% of the water volume and is relatively colder, saltier, denser, and nutrient rich.

Surface waters circulate about the ocean in great basin-wide circles called gyres, which are driven by winds. Because water has a high capacity for storing heat, this process transports heat from the tropics, where solar heating is intense, to the cooler poles. A well-known example is the Gulf Stream that travels from the Gulf of Mexico northeast to Europe, thereby contributing to the fact that Europe's temperature is higherthan would be expected given its latitude. During the winter, the density of seawater increases as it approaches the poles due to loss of heat and freshwater; freshwater freezes out onto the polar ice caps leaving behind salts. The increase in density is sufficient to cause surface water to sink into a deep-water layer of the same density. The sinking water displaces existing water, triggering a deep-water current that eventually rises back to the surface at tropical latitudes. The deepest ocean currents may take hundreds to thousands of years to reach the surface again.

The long circulation time of deep ocean currents hold an important lesson for those interested in the effects of global warming. Once the chemistry or heat content of the deep ocean changes, it will remain so much longer than a change in the atmosphere or ocean surface. Gases and heat absorbed by the ocean now will eventually be released again to the atmosphere. Thus, we cannot expect to quickly reverse the course of greenhouse warming once it begins.

The longest time-scale relevant to element cycling is the circulation of rocks. The crust and upper mantle form a relatively rigid layer at Earth's surface, called the lithosphere, that ranges from 50 to 200 km thick. The lithosphere is divided into eight major plates that fit together like a puzzle to form the patchwork of continents and ocean basins we have today. Below the lithosphere is the asthenosphere (also part of the mantle) that is relatively plastic and fluid. Convection currents in the asthenosphere cause the lithospheric plates to move about the surface, at various times sliding past one another, pulling apart or colliding.

New ocean crust is created when molten rock (magma) from the mantle rises up to fill cracks created where plates separate. The hot magma crystallizes to form igneous rocks (Figure 2). One of the most active ocean-spreading regions is the mid-Atlantic ridge, home to a unique marine ecosystem that derives energy from reduced chemicals in sea water emanating from the sea floor. Plates remelt when they are forced back down into the mantle during collisions in a process called subduction. Subduction occurs whenever the leading edge of one of the plates bears oceanic crust. If one plate bears the relatively light rocks of the continental crust, the oceanic crust is always subducted. If both plates bear oceanic crust, then one will be subducted into the mantle and the other will remain at the surface. Oceanic crust circulates through this pathway every 110 10⁶ y170 10⁶ y, carrying elements along for the journey.

Figure 2. A generalized diagram of the steady-state rock cycle
Sediments and continental crystalline crust masses are in units of metric tons (1 10⁶ g), and fluxes are metric tons per year. Total sedimentation is 9 10⁹ metric tons annually. Source: Gregor C.B. (1988). Prologue, cyclic processes in geology, a historical sketch. Chemical Cycles and the Evolution of the Earth (ed. C.B. Gregor, R.M. Garrels, F.T. Mackenzie, and J.G. Maynard), pp. 516. New York: John Wiley and Sons. Copyright 1988 Wiley. Reprinted with permission by John Wiley & Sons, Inc.

A somewhat slower rock cycle begins when two plates bearing continental crust collide. Neither plate is subducted, but rather both are crushed and deformed, often folding to form new mountain ranges. Over geologic time, the mountains erode away and are carried to oceans as sediments and dissolved ions. Ocean sediments and salts undergo diagenesis to sedimentary rocks as they are buried. Under extreme pressure and heat, the crystal structure of sedimentary rocks changes creating metamorphic rocks.

Elements bound into rocks are released in the process of rock weathering. In physical weathering the rock is broken into smaller and smaller pieces, but the chemical composition of the rock is largely unchanged. The primary agents of physical weathering are water and roots. Liquid water seeps into cracks in the rock then expands upon freezing; plant roots operate in a similar manner by slowly expanding during growth.

In chemical weathering the elemental composition of the rock changes as particular elements are released and others stay behind. The elements that are released can reform into secondary (clay) minerals such as kaolinite. This process is greatly enhanced by the activity of plant roots and soil microbes that produce CO₂ during respiration. The CO₂ dissolves in water forming carbonic acid.

CO₂ + H₂O H₂CO₃ H⁺ + HCO₃                             (1)

Carbonic acid, warm temperatures, and low pH all enhance the rate of chemical weathering. This is a key feature in the cycle of many elements for which rocks are their ultimate source, including most nutrients. The group includes phosphorus (P), calcium (Ca), sodium (Na), potassium (K), magnesium (Mg), iron (Fe), manganese (Mn), zinc (Zn), and several more trace elements. Because these elements are essential nutrients, the rate of rock weathering places a long-term limit on the potential productivity of ecosystems (see Clay Minerals in Mineralogy).

A number of the environmental problems that we face today are caused by a human-induced increase in rock "weathering" through mining activity. For example, mining provides phosphorus, a key ingredient in fertilizers that greatly enhances the yield of agricultural crops. Unfortunately, phosphorus moves in water from agricultural fields to streams, lakes, and estuaries where it often overstimulates aquatic productivity. Excess phosphorus may cause changes that range from shifts in community composition, to losses of species diversity and fish mortality. Mining has also increased the abundance of lead (Pb), zinc (Zn), mercury (Hg), selenium (Se), and other metals that are toxic even at low concentrations.

An important product of the rock weathering process is soil. Soil is the biologically active layer of Earth's crust and is composed of minerals, organic matter, and biota. Roughly 50% of a soil's volume is solids (minerals and organic matter) and 50% is space (air and water). Soil solids act as both an element source for organisms and an element sink. Weathering releases elements from minerals, while decomposition (i.e., mineralization) releases elements from organic matter. Many important characteristics of soils are derived from the proportion of sand, silt, and clay minerals, which differ in size and charge. The largest particles are sand (0.052.0 mm diameter), followed by silt (0.0020.05 mm), and then clay (<0.002 mm). The amount of clay is a particularly important feature of soils because clay particles carry a negative charge that forms electrostatic bonds with nutrient cations such as NH₄⁺, Ca²⁺, and K⁺. Soil organic matter is also negatively charged and serves the same nutrient storage function.

The soil atmosphere is a much more dynamic environment than the free atmosphere in which people live. Soil O₂ levels vary from 21% (the same as air) under relatively dry conditions to 0% after soil pores have been saturated with water for a period of days or weeks. Wetland plants have physiological and morphological adaptations that allow them to continue aerobic respiration even when flooded, but the absence of O₂ has profound changes on microbial communities, which shift from dominantly aerobic to anaerobic species (see Soils).

The most important cogs in any nutrient cycle are: (1) the producers that combine inorganic elements and light (or chemical energy) into organic compounds, (2) the decomposers that convert the organically bound elements back to an inorganic form, and (3) the air and water that transports the elements within and between ecosystems. In fact, the simplest ecosystem that could theoretically exist would consist of one producer and one decomposer.

Life is a dominant factor regulating element cycles. Organisms exploit the unique properties of elements for energy and nutrients. Not surprisingly, life has evolved the biochemical machinery to enhance rates of the chemical transformations on which they depend. For example, organisms increase the supply of certain forms of nitrogen and phosphorus by accelerating nitrogen fixation, rock weathering, and detritus decomposition. Biota also influence nutrient cycles that were previously considered to be largely abiotic. For example, biota influence nutrient inputs from precipitation, dust deposition, and rock weathering; they influence nutrient exports by mineral precipitation, erosion, stream flow, flooding, and leaching to groundwater. Because organisms and chemical cycling are intimately coupled in ecosystems, neither can be understood in isolation from the other. Reciprocal interactions between organisms and their chemical environment present some of the most basic questions facing scientists:

To what extent do element cycles regulate the distribution and abundance of organisms?

To what extent does the distribution and abundance of organisms regulate element cycling?

Does high biological diversity help stabilize element cycles?

Do feedbacks exist between biota and climate that either dampen or amplify climate change?

The carbon cycle is an ideal example of how biological, chemical, and physical processes interact to influence the capacity of Earth to support life. Carbon is a versatile element. It has the unusual property of bonding to itself (i.e., CC bonds) to form short chains, long chains, branching chains, and rings. This versatility in structure begets versatility in function, which ranges from light capture to information storage. The role of carbon in energy storage and transformation is due to its ability to exist in several oxidation states. As gases, most carbon compounds adsorb strongly in the infrared region of the electromagnetic spectrum, which makes them powerful heat-trapping (greenhouse) gases. These features are all absent in silica, another abundant element and a neighbor of carbon in group IVA of the periodic chart. Carbon accounts for about 50% of all biomass and seems particularly suited to support life.

Plants, phytoplankton, and other autotrophic organisms consume about 30% of the all the CO₂ in the atmosphere each year. The vast majority of CO₂ assimilation is due to photosynthesis, which is driven by sunlight at wavelengths of 400700 nm (i.e., photosynthetically active radiation). The light energy splits water into oxygen (O), protons (H⁺), and electrons (e):

sunlight + 12H₂O 6O₂ + 24e + 24H⁺                     (2)

This initial reaction provides a source of reducing power to drive the reduction of CO₂:

6CO₂ + 24e + 24H⁺ C₆H₁₂O₆ + 6H₂O                         (3)

In this example, the combined reaction yields glucose, a compound that is energy rich by virtue of six reduced carbon atoms:

6CO₂ + 6H₂O C₆H₁₂O₆ + 6O₂                         (4)

The amount of CO₂ assimilated through photosynthesis by a plant, a forest, or the globe over a given period of time is its gross primary productivity (GPP). The gross primary production of terrestrial vegetation globally is roughly 120 10¹⁵ g y¹ (Figure 3).

Figure 3. The present-day carbon cycle. All pools are expressed in units of 1 10¹⁵ gC and all fluxes in units of 1 10¹⁵ gC y¹, averaged for the 1980s. Source: Schlesinger W.H. (1997). Biogeochemistry: An Analysis of Global Change, 443 pp. San Diego: Academic Press.

About 50% of gross primary production is used by the plant to supply energy for the construction of new tissues and maintenance of existing tissues. The energy for this activity is supplied by oxidizing or burning organic carbon substrates, such as carbohydrates, in a processes known as respiration. A plant can grow only if gross primary production exceeds respiration, otherwise it will lack the structural carbon required to to build cellulose and other new tissues. Excess carbon is combined with N, P, and many other elements to build new roots and shoots. This new growth is called net primary production (NPP), the quantity we recognize as plant yield. The net primary production of terrestrial vegetation is about 60 10¹⁵ g y¹ (Figure 3).

At a global scale, net primary production is controlled primarily by temperature and precipitation. The highest terrestrial net primary production is in tropical rainforests that are warm and wet year-round, while the lowest is in arctic tundra and deserts. Within a given ecosystem type, soil fertility and local environmental factors such as microclimate and flooding exert strong controls on productivity.

Nutrient availability is a key factor controlling the productivity of oceans. A disproportionately large fraction of the productivity occurs near continents where nutrient-laden rivers empty into the ocean. Productivity is also high in upwelling zones because of dissolved nutrients that have accumulated in deep ocean water from decomposition of organic matter.

Human populations have grown to the point that we now influence the fate of about 40% of Earths terrestrial net primary productivity. Peter Vitousek and colleagues estimate that humans use roughly 2% of terrestrial net primary productivity by consuming wood, livestock, and crops directly. The figure is much larger if one appreciates the fact that all the net primary productivity on managed lands is influenced by human activity, not just the portion consumed directly. Thus, the full net primary production of natural systems has been co-opted wherever humans have converted them into pastures, crop farms, tree farms, or urban areas. Human use of freshwater and marine net primary productivity is somewhat lower, yielding a combined estimate of 25% for terrestrial and aquatic systems.

Aerobic respiration is essentially the reverse of photosynthesis, although energy is released as adenosine triphosphate (ATP) rather than sunlight:

C₆H₁₂O₂ + 6H₂O 6CO₂ + 24e + 24H⁺                     (5)

24e+24H⁺ 6O₂ 6H₂O + 36ATP                     (6)

The combined reaction yields adenosine triphosphate (ATP), which contributes chemical energy to most metabolic reactions in organisms.

Perhaps the most significant component in respiration is the tiny electron (e) This is because the energy in ATP is derived from chemical gradients created by an electron transport chain. Respiration will not occur without both an electron source and an electron sink. In aerobic respiration, O₂ serves as the electron sink. Respiration can occur in the absence of O₂ provided there is an alternative electron acceptor available.

There are a wide variety of electron acceptors used by microorganisms that inhabit anaerobic environments including nitrate (NO₃), ferric iron (Fe³⁺), sulfate (SO₄²), and carbon dioxide in the form of bicarbonate ion (HCO₃). However, all these organisms use the same electron source: detrital carbon. Carbon availability often limits microbial activity in ecosystems. Competition for carbon substrates tends to favor those organisms using the electron acceptors that yield the most energy: O₂ > NO₃ > Fe³⁺ > SO₄^2- > CO₂ (Table 1). One group of microbes dominates carbon consumption until its electron acceptor is exhausted, followed by the next most favorable pathway, and so on. In the sediments of oceans, mangroves, and salt marshes, sulfate-reducing bacteria dominate carbon anaerobic metabolism because sulfate reduction is favored energetically over CO₂-reduction (i.e., methanogenesis), and SO₄^2- is far more abundant than NO₃. In freshwater ecosystems, SO₄^2- concentrations are quite low and methanogenesis is a far more dominant microbial decomposition pathway.

Table 1. Thermodynamic sequence for reduction of inorganic substances by organic matter

Heterotrophic organisms fuel respiration with organic carbon compounds that were ultimately produced by plants or other autotrophs. In many cases, plant biomass is also the ultimate source of nutrients that heterotrophs require. In the process of extracting energy and nutrients, such organisms cause organic compounds to decompose and thereby provide raw materials for new productivity. Decomposition vastly increases the productivity of Earth by preventing nutrient elements from being locked away indefinitely in soils and sediments.

Decomposition proceeds by both physical and chemical means. Earthworms, microarthropods, and other animals are effective at breaking large pieces of organic matter, such as leaves and stems, into smaller pieces with greater surface area that are more vulnerable to chemical attack. Bacteria, fungi, and plants secrete extracellular enzymes that selectively cleave covalent bonds and ultimately convert organically bound elements into inorganic forms such as CO₂, NH₄⁺, and PO₄³. This process of mineralization often supplies most of the annual nutrient requirement for biological productivity.

Terrestrial plants provide about half of the global detritus production that supports decomposers. Plant detritus is composed primarily of proteins, holocellulose, and lignin. Holocellulose includes both cellulose and hemicellulose and it is the single most abundant form of carbon in woody and nonwoody plant tissues. Leaves have higher concentrations of protein than wood because of the abundance of photosynthetic enzymes such as ribulose bisphosphate carboxylase (i.e., Rubisco), the enzyme that converts CO₂ to sugar in the first step of photosynthesis. Both holocellulose and proteins are rapidly decomposed by microorganisms. Lignin is most abundant in wood and is highly resistant to decomposition.

Most chemical decomposition is mediated by bacteria or fungi with enzymes such as cellulase, pectinase, and lignin peroxidase, which selectively decompose cellulose, pectin, or lignin. Both groups attack cellulose, but fungi are responsible for most lignin decomposition. Termites are able to consume wood because they harbor streptomycetes fungi in their foregut. Ants belonging to the tropical group Attinae actively cultivate fungi on leaves in underground nests, then consume the fungus directly. Mutualistic relationships between structurally complex organisms such as plants and animals, and biochemically sophisticated microorganisms are a common theme in nature and serve to illustrate why microorganisms are a key biological cog in any biogeochemical cycle.

The rate of organic matter decomposition is generally highest immediately following death and then decreases with time. In fact, decomposition is often modeled as a simple exponential decay curve

where M₀ is the initial mass, M is the mass that remains after a period of time (t), k is a decay constant (units = 1/t). Notice that k is negative because mass is lost during decomposition. The exponential shape of a decay curve reflects differences in the behavior of the various chemical constituents discussed above. The mass of a leaf is largely proteins and cellulose, which are easily degraded by microorganisms, and therefore disappear quickly from the detritus. Decomposition slows considerably as more recalcitrant compounds (e.g., lignin) come to dominate the remaining mass.

The two main environmental factors that govern decomposition at a global scale are the same ones that govern primary productivitytemperature and precipitation. In fact, decomposition rates typically double with a 10 C increase in temperature. A third key factor is the chemical composition of the detritus, sometimes referred to as the carbon quality. Detritus is generally poor in nitrogen and rich in carbon compared to the nutritional needs of decomposing fungi and bacteria. Detritus is considered to be of high quality (i.e. easily decomposable) if it relatively high in nitrogen and low in carbon-dominated compounds such as lignin. Indices of carbon quality that have proven exceptionally useful are C-to-N ratio, the lignin content, and the lignin-to-N ratio. Leaves have low values for each of these indices and thus decompose rapidly, while wood has high values and decomposes more slowly. See Decomposition.

Most of the gross primary productivity of the planet is consumed each year in the respiration of plants, animals and microorganisms. The portion that remains behind is the net ecosystem production (NEP)

NEP = GPP R_P R_H                         (8)

where GPP is gross primary production, R_P is plant respiration and R_H is heterotrophic respiration. Although net ecosystem production is usually a small (<1%) proportion of the annual GPP, over time this flux is responsible for globally significant pools of carbon stored in wood, soils, and aquatic sediments. Because the carbon is in an organic form, NEP is also a long-term sink for N, P, Ca, and nutrient elements.

Rates of net ecosystem production are highest in ecosystems with high productivity and a large capacity to sequester (i.e. store) carbon. When considered over annual to decadal time scales, forests have a high capacity to sequester carbon as wood. Sequestration in upland soils is a slower process, but it can be fairly rapid in the anaerobic sediments of wetlands, lakes, and the continental shelf. The absence of O₂ significantly slows the pace of microbial decomposition, even though electron acceptors other than O₂ are available. Cold temperatures also inhibits decomposition. For these reasons, the vast peatland complexes (bogs and fens) of the boreal zone hold at least 25% of the terrestrial soil carbon globally.

Notice that net ecosystem production is different from related terms introduced so far because it is a quantity that can be either positive or negative. NEP is positive when there is a net transfer of carbon from the atmosphere to the ecosystem; it is negative when there is a net transfer of carbon from the ecosystem to the atmosphere. A major interest of scientists working on the issue of global climate change is whether elevated CO₂ and associated greenhouse gas-induced warming will increase or decrease NEP. In particular, there is concern that some ecosystems, such as boreal peatlands, may become drier and switch from net carbon sinks (+NEP) to net carbon sources (NEP) with respect to the atmosphere. Fires will play an important role in determining carbon balances because they can quickly mobilize large amounts of accumulated NEP back into the atmosphere. Feedbacks between climate and biogeochemical cycles have the potential to influence the course of climate change during this century.

A large portion of the carbon stored in ocean sediments is in the form of carbonates (e.g., CaCO₃). Large amounts of CaCO₃ are produced annually by marine organisms such as clams, oysters, foraminifera, coccolithophores, and pteropods that actively incorporate the material into their shells. This biogenic reaction consumes bicarbonate (HCO₃) and calcium ions (Ca²⁺):

Ca²⁺ + 2HCO₃ CaCO₃ + H₂O + CO₂ (9)

Much of the carbonate production sinks through the water column and is stored in shallow marine sediments (<4500 m). At depths >4500 m, the carbonate is driven back into solution by the higher pressure and lower pH characteristics of deep water. See Carbonates in Sedimentary Rocks.

As in oceans, there are terrestrial processes that sequester carbon in an inorganic form. The soils of arid and semiarid regions hold 800 10¹⁵ gC as CaCO₃ globally, but the annual rate of formation is rather slow. Soil and microbial respiration typically raise soil CO₂ concentrations to levels of 13%, or about two orders of magnitude above atmospheric concentrations. This biotic process increases the inorganic carbon content of groundwater, generating a long-term carbon sink lasting hundreds to thousands of years.

Organisms require 23 elements for energy metabolism and biomass production. The elements that most often limit productivitynitrogen and phosphorusare both in great demand by organisms and somewhat unavailable in the environment. Elements that organisms need in relatively large amounts also include carbon, sulfur, oxygen, hydrogen, potassium, calcium, iron, and magnesium (see Biogeochemical Cycling of Micronutrients). Collectively, these elements constitute the macronutrients. The remaining elements, micronutrients, are required in trace amounts but can nonetheless limit primary production (Table 2).

Table 2. Concentrations of macronutrients and micronutrients in leaves that are considered deficient, adequate, or toxic for various species of crop plants. Nonessential elements are listed at concentrations typically found in crops.

Carbon is linked to other element cycles through the element stoichiometry of life, meaning the ratios of carbon to nitrogen, phosphorus, and other essential nutrients. An influential application of this concept was proposed in 1958 by Albert Redfield, who observed that element ratios in marine plankton are fairly uniform at a mean chemical composition of 106 carbon atoms to 16 nitrogen atoms to 1 phosphorus atom. The Redfield ratio of 106C:16N:1P can be used to calculate the nutrient requirement to support a given level of marine productivity. For example, provided the net primary production of oceans is about 4.2 10¹⁵ moles C y¹ (i.e., 50 10¹⁵ g C y¹), the nitrogen required to support ocean productivity is 0.6 10¹⁵ moles N y^-1. Just 14% of this amount is delivered to oceans by rivers, upwelling, and the atmosphere, suggesting that 86% is derived from internal recycling!

The element ratios of terrestrial plants tend to be more variable than plankton because of their structural tissues. Plants require nitrogen in combination with carbon in ratios that range from about 30:1 (C:N) in foliage to 160:1 in wood. Microorganisms have a somewhat higher nitrogen requirement with C:N ratios of about 12:1. Because of the disparity between the nutrient content of plant and microbial biomass, decomposers compete with plants for nitrogen and phosphorus from the soil. In fact, nitrogen usually accumulates in decomposing detritus for a time until the C:N ratio of detritus begins to approach that of microbes.

The Spengel-Leibig law of the minimum states that, under conditions of sufficient light and temperature, plant growth is limited by the single nutrient element for which the imbalance between supply and demand is greatest. The law applies to both terrestrial and aquatic ecosystems and it explains why nitrogen or phosphorus pollution can cause excessive aquatic productivity or eutrophication. Ratios of bicarbonate (HCO₃), nitrate (NO₃), and phosphate (HPO₄²) in seawater are about 1017:15:1. Comparing these ratios to the Redfield ratio suggests that ocean production may be slightly nitrogen limited, but certainly not carbon limited.

Nitrogen limits the productivity of most terrestrial and marine ecosystems, despite the fact that it constitutes 78% of Earth's atmosphere in the form of N₂ (dinitrogen gas). In one of the great ironies of nature, no plant or animal can directly use N₂ for a nitrogen source even though they are bathed in the gas. Rather, plants take nitrogen from the soil as NH₄⁺ (ammonium), NO₃ (nitrate), or organic-N, and animals eat plants (or other animals). Because rocks contain little to no nitrogen, the atmosphere is the ultimate source of nitrogen in soils and sediments. Atmospheric N₂ is converted to the biologically available forms in soils during the process of nitrogen fixation.

Before the evolution of nitrogen-fixing organisms, most N₂ was fixed during lightning strikes. Indeed, a great deal of energy is required to break the covalent triple bond in N₂, a task that most organisms cannot perform. A key exception is N₂-fixing bacteria that assimilate N₂ into organic-nitrogen compounds. Some plants are able to grow well on nutrient deficient soils because they evolved mutualistic relationships with N₂-fixing bacteria. One such group is the legume plant family that includes peas, soybeans, and most other beans. Legumes provide organic carbon (energy) to N₂-fixing bacteria, and receive in return a useable form of nitrogen. The symbiotic N₂-fixing bacteria, such as Rhizobium spp., live in spherical nodules on the plant root. N₂-fixing bacteria contribute about 70% of the current annual N₂ fixation, human production of fertilizers is roughly 25%, and lightning discharge is 5% (Figure 4).

Figure 4. The nitrogen cycle. All pools are expressed in units of 1 10¹² g N and all fluxes in units of 1 10¹² g N y¹. Source: Schlesinger, W.H. (1997). Biogeochemistry: An analysis of global change, 443 pp. San Diego: Academic Press.

Relatively little of the nitrogen required for the productivity of ecosystems comes from "new" nitrogen inputs via N₂-fixation. Rather, most nitrogen is supplied by decomposition, also called mineralization, which yields ammonium (NH₄⁺) and organic-nitrogen compounds (Figure 5). Decomposition represents an internal recycling system. Plants also recycle nitrogen directly by absorbing it from dying leaves and stems before they fall. Many people who live in the temperate zone enjoy watching this process in the autumn when deciduous trees turn from green to red, orange, or yellowa fine example of the beauty in biogeochemistry! The plant is removing nitrogen from chlorophyll and other nitrogen-rich pigments and enzymes in the leaf, and storing it in stems and roots. The plant then uses the stored nitrogen to build new foliage in the following spring.

Figure 5. Key microbial processes in the nitrogen cycle. The gases NO (nitric oxide) and N₂O (nitrous oxide) are produced when a portion of the ammonium pool undergoing nitrification is not fully oxidized; similarly these trace gases are produced from incomplete reduction during denitrification. Source: Firestone M.K. and Davidson E.A. (1989). Microbiological basis of NO and N₂O production and consumption in soil. Exchange of Trace Gases Between Terrestrial Ecosystems and the Atmosphere (ed. M.O. Andreae and D.S. Schimel), pp. 721. New York: John Wiley and Sons.

Nitrogen exists in several oxidation states, a fact that microbes have exploited to derive metabolic energy. The oxidation state of both nitrogen atoms in N₂ is zero. Mineralization to NH₄⁺ yields N(III) (nitrogen with an oxidation state of 3). Under aerobic conditions, NH₄⁺ is oxidized by nitrifying bacteria to NO₃ (nitrate, N(+V)) in the two-step process of nitrification (Figure 5). In the first step, NH₄⁺ is oxidized to NO₂ (nitrite, N(+III)) by autotrophic bacteria such as Nitrosomonas species, which use the energy to fix CO₂ directly from the atmosphere. The second step is performed by autotrophic bacteria such as Nitrobacter species that oxidize NO₂ to NO₃.

If N₂ fixation was the only process operating on the N₂ pool of the atmosphere it would have long since depleted the atmosphere entirely. However, under anaerobic conditions nitrogen is returned to the atmosphere by the process of denitrification. Denitrification contributes to the N limitation of aquatic and terrestrial ecosystems, and it is a key process for removing NO₃ from polluted rivers and groundwater supplies. Wetlands are valuable features in human-dominated landscapes, in part because they support rapid active denitrification of polluted water. The process also occurs in the anaerobic microsites of upland soils. Under certain conditions, a portion of the nitrogen subjected to nitrification and denitrification is lost to the atmosphere as NO (nitric oxide) or N₂O (nitrous oxide). Nitric oxide is a reactive gas that participates in several key atmospheric reactions that consume or produce ozone. Nitrous oxide is relatively unreactive, but it is a powerful greenhouse gas. Clearly, long-term nitrogen availability in ecosystems is determined by the balance between inputs of nitrogen (e.g., N₂-fixation and atmospheric deposition) and outputs (denitrification and hydrologic losses).

The phosphorus cycle is an interesting contrast to the nitrogen cycle because it does not have an important gas phase nor does it readily undergo reductionoxidation reactions. Rocks are the ultimate source of the phosphorus in ecosystems (Figure 6). It is bound in minerals such as apatite (Ca₃(PO₄)₂) and released upon rock weathering, a slow process in which minerals physically and chemically degrade into their constituent elements.

Figure 6. The phosphorus cycle .All pools are expressed in units of 1 10¹² g P and all fluxes in units of 1 10¹² g P y^-1. Source: Schlesinger W.H. (1997). Biogeochemistry: an analysis of global change, 443 pp. San Diego: Academic Press.

The phosphorus and nitrogen cycles are also similar is some respects. They are the two most limiting nutrients in the biosphere. Most of the phosphorus demand of ecosystems is supplied through mineralization of organic matter rather than "new" inputs. Plants have mutualistic relationships with mycorrhizal fungi that greatly enhance phosphorus acquisition, and, to a lesser extent, nitrogen. Humans have increased rates of phosphorus input from weathering through mining.

Because nitrogen is made available by biological processes (i.e., N₂ fixation), it should never be in short supply compared to phosphorus, for which there is no biological equivalent to N₂ fixation. Thus, one can ask why is nitrogen ever limiting? It has been found that phosphorus typically limits primary production in freshwater ecosystems, due in large part to the presence of N₂-fixing organisms. However, nitrogen limits primary production in most terrestrial and coastal marine ecosystems, including salt marshes.

Several processes may contribute to persistent nitrogen limitation. Phosphorus is retained in soils and sediments through biological uptake and the formation of insoluble compounds with common cations such as Ca, Fe, and Al. The nitrogen cycle lacks these geochemical retention mechanisms. Unlike phosphorus, the nitrogen cycle has a dominant gas phase and can be lost to the atmosphere as N₂, N₂O, and NO during microbial transformations or as a result of fire. Another possibility is that N₂-fixing organisms are themselves limited by the availability of phosphorus, iron, molybdenum or other nutrients; thus, what appears to be nitrogen limitation is actually limitation by some other nutrient in disguise.

Soil age may sometimes determine if a terrestrial ecosystem is nitrogen or phosphorus limited. Young soils lack nitrogen because of the short history of N₂-fixing organisms growing on them, but can have abundant phosphorus, which comes from rocks. As soils age, nitrogen accumulates from biological N₂-fixation, while phosphorus availability declines due to fixation in secondary minerals. This pattern has been elucidated in Hawaiian forests with a series of highly controlled fertilization studies. However, other studies suggest nitrogen limitation persists in temperate and boreal forests that have had 12 000 years to 20 000 years to accumulate nitrogen, and nitrogen limitation prevails in some aquatic systems that are old enough to have accumulated a great deal of nitrogen through biological N₂ fixation.

There are a number of nutrients besides nitrogen and phosphorus that are required by organisms in relatively large amounts. Sulfur has a complex cycle that includes several oxidationreduction states; gas, liquid, and solid phases; biological production and consumption; and strong interactions with metal cations. As discussed earlier, sulfate (SO₄²) reduction is the dominant pathway of carbon mineralization in marine and salt marsh ecosystems. It produces H₂S (hydrogen sulfide), which has an unpleasant rotten-egg odor familiar to anyone who has tromped through a saltmarsh. H₂S reacts with Fe(III) to produce iron sulfide (FeS) and pyrite (FeS₂), also known as fools gold. When exposed to O₂, pyrite-bound sulfur oxidizes to SO₄², producing acidity. Acid mine drainage causes extremely low pH levels (<3) wherever streams drain through abandoned mineral or coal mines. Sulfidic ores also contain trace metals that are released during pyrite oxidation, further exacerbating the environmental impacts of mining. Dimethyl sulfide is a sulfur-bearing gas emitted by marine algae that oxidizes in the atmosphere to produce radiation-scattering sulfate aerosols. Sulfate aerosols cool the planet by reflecting solar radiation back to space, thereby offsetting the greenhouse warming caused by CO₂, CH₄ and other gases. This form of sulfur is notable because it has been implicated as a mechanism by which life on Earth can modify climate.

In addition to P, a number of other important macronutrients are derived from rock weathering including Ca (calcium), Mg (magnesium), Na (sodium), and K (potassium). Under certain conditions, each of these nutrients can severely limit primary productivity. Mammals can often be limited by Na because their demand is much higher than typical levels found in soils and plants. K is lost quickly from soils through leaching in groundwater (it is not strongly adsorbed by soils), and through incorporation into clay minerals. For this reason it is a common component in commercial fertilizers.

Many elements are required by organisms in relatively smaller amounts than the macronutrients. In some cases, these elements are abundant in the environment, but tend to occur in forms that are difficult for organisms to access. One example is iron (Fe).

The iron cycle exerts a strong influence on the biogeochemistry of several other elements in soils including C, P, S, Mn, and many trace metals. Under anaerobic conditions, microbial reduction of Fe(III) is coupled to carbon oxidation, a process that competes with methanogenesis and SO₄ reduction and for organic carbon substrates. Fe reduction consumes between 21% and 100% of the organic carbon in some anaerobic sediments. In an oxidized state, Fe is a rust-colored solid of the general form Fe(OH)₃. But in a reduced form (i.e., Fe(II)), it is water soluble. Where aerobic and anaerobic environments meet, Fe(II) is oxidized back to Fe(III) in a process that can be mediated by Fe-oxidizing microbes:

Fe²⁺ + O₂ + H⁺ Fe³⁺ + H₂O                     (10)

Fe³⁺ + 3H₂O Fe(OH)₃ + 3H⁺                         (11)

This pattern, in which an element is alternately reduced and oxidized across an anaerobicaerobic interface, is an important feature in the cycles of several elements. Examples include CH₄ (CH₄ CO₂) and H₂S (H₂S SO₄²). Various microbes mediate the two sides of these reductionoxidation reactions, deriving energy at each step.

Iron is required by many key enzyme systems. Some aerobic bacteria and fungi produce strong chelating agents called siderophores, which are released into the soil where they bind and adsorb Fe(III). Plants produce similar compounds known as phytosiderophores.

Other metal ions that undergo reductionoxidation reactions include Mn (manganese) and Cu (copper). Some marine animals require relatively large amounts of Cu for a protein called hemocyanin, which is similar in structure and function to the O₂-carrying protein hemoglobin. Zinc (Zn) differs from Fe and Cu because it does not change its oxidation state and exists only as Zn(II). It readily forms sulfide ores such as wurzite (ZnS).

5.6. Contaminants

Trace elements are 69 of the 79 naturally occurring elements in Earth's crust, yet they total <0.5% of the crust by mass. By definition, their individual abundance is <0.1% of crustal mass. This definition is somewhat misleading because certain trace elements can occur locally at high concentrations (e.g., gold nuggets). Many trace elements are both nutrients and contaminants depending on their concentration (Table 2). At low concentrations they may limit productivity because demand exceeds supply, while at high concentrations they suppress productivity because they produce toxic physiological effects. As a group, their ecological and economic importance is far greater than their abundance would suggest.

Trace element cycles begin with rock weathering, but subsequent steps include solution phases, gas phases, and biological transformations. Most are sensitive to environmental factors such as reductionoxidation potential, pH, and the concentration of reactive compounds such as H₂S. A major new force in the cycles of all trace elements is human activity. Mobilization of trace elements through mining and the manufacture of novel compounds have increased the background concentrations of these chemicals in many parts of the world. For example, mining of copper (Cu), lead (Pb), nickel (Ni), tin (Sn), and zinc (Zn) have increased eightfold, while human population doubled, in the latter half of the twentieth century. The increase in trace elements loads in ecosystems is due both to human population growth and increasing per-capita demand for resources.

Major anthropogenic sources of trace elements include coal combustion, fossil fuel combustion, and the chemical production industry. The chemical industry also produces many novel compounds with uncertain effects on natural ecosystems and human health. By one estimate, the industry produces 80 000 different chemicals. Among the most toxic of elements are arsenic (Ar), cadmium (Cd), mercury (Hg), and lead (Pb). Each of these elements form compounds with sulfur (metal sulfides), adsorbs strongly to soils, and is transported by both water and air.

Arsenic is a poison used to control insects and rodents. It is fairly volatile and readily enters the atmosphere as As₂O₃ during fossil-fuel combustion. Concentrations in drinking water >50 m g L^-1 are usually considered unsafe. Cadmium is highly toxic to humans and has been implicated in a disease of the bones called "itai-itai" in Japan. Only 15% of the atmospheric deposition is from natural sources. It is relatively mobile in aquatic systems, and tends to bind well to sediments and organic matter. Mercury has many chemical forms and is highly soluble when methylated (i.e., HgCH₃). Human-related Hg emissions to the atmosphere are growing 35% per year in developing countries. Lead is moderately toxic to plants and animals. About 96% of lead emissions are anthropogenic sources such as metal manufacture, coal combustion, cement production, and sewage sludge disposal. Like mercury, lead is biologically methylated in anaerobic sediments, producing a gaseous form that can be widely distributed in the environment. Lead emissions declined substantially once its use as a fossil fuel additive was phased out beginning in the mid-1970s.

Background levels of organic contaminants are negligible, and the major source is human activity. Pesticides such as DDT (dichlorodiphenyltrichloroethane) are widely used to control insects, rodents, weeds, and fungi, but they are also widely known to cause health problems in nontarget wildlife and humans. Other chemicals produced by humans that are particularly toxic include polychlorinated biphenyls (PCBs), polyaromatic hydrocarbons (PAHs), benzene and related compounds, and dioxins.

An important feature common to trace elements and contaminants is their propensity to increase in concentration as they are passed up the food chain. This process is called biomagnification and it accounts in part for the toxic impacts of these chemicals in ecosystems. Important sinks for contaminants include sediments and microbial degradation. The use of microbes and plants to clean up toxic waste has proven to be a successful strategy. See Toxic Waste.

More than 99.9% of Earths atmosphere is composed of just three gasesnitrogen (N₂), oxygen (O₂), and argon (Ar; Table 3). The remainder of the gases occur at trace levels, but nonetheless exert a powerful influence on climate by trapping infrared energy (heat) near the planets surface. Such gases include water vapor (H₂O), carbon dioxide (CO₂), methane (CH₄), nitrous oxide (N₂O), ozone (O₃), and a variety of carbon compounds containing chlorine and fluorine (i.e., CFCs or chlofluorocarbons). These gases raise the global surface temperatures a full 30 C to the present mean of about 15 C. Without the natural greenhouse effect much of the biosphere would be frozen!

Table 3. Composition of gases in Earths atmosphere and their potential to cause global warming. Notice the vast difference in concentration among the gases.

Water vapor is the most abundant greenhouse gas and contributes most of the natural greenhouse effect. Because H₂O vapor can rapidly precipitate to form a liquid or solid, atmospheric concentrations respond strongly to changes in temperature initiated by other causes. In many cases, water vapor levels increase in step with increasing temperature, thereby causing further increases in temperaturea positive feedback loop. Negative feedbacks are also possible because some types of clouds have a net cooling influence on climate.

Anthropogenic activity has increased concentrations of a number of other greenhouse gases, most notably CO₂, CH₄, and N₂O (see section 7 below). The cycling of CO₂ was discussed in detail above (see section 4). Here we will review some details of CH₄ and N₂O cycling.

Methane is an odorless gas that is increasing in the atmosphere at about 1% annually. It is produced through geological processes operating on organic carbon in marine sediments, but most of the annual production is a by-product of the metabolism by methanogenic bacteria. Methanogenic metabolism evolved well before O₂-producing photosynthesis, but these organisms are now restricted to anaerobic environments such as wetland soils, aquatic sediments, aquifers and the intestinal tracts of animals. Natural environments such as wetlands, aquatic sediments, and termites were once the major sources of CH₄ emitted to the atmosphere. At the end of the twentieth century, about 30% comes from natural sources and the remainder from anthropogenic sources such as landfills, coal mines, natural gas mining, animal wastes, cattle, and rice paddies.

Methanogenesis is the final step in a process that ultimately begins with photosynthetic assimilation of CO₂ into organic compounds by plants, algae, and other autotrophic organisms. These compounds supply the energy required for heterotrophic metabolism. Because of competition for electron donors, high rates of methanogenesis are confined to areas where sulfate reduction is limited by the SO₄^2- supply, a common condition in freshwater sediments.

Detrital organic carbon undergoes a series of fermentation reactions, each mediated by a different population of bacteria. Methanogens can use a wide variety of simple organic carbon compounds, but the two most common pathways are CO₂ reduction,

CO₂ + 4H₂ CH₄ + 2H₂O                     (12)

and acetate fermentation,

CH₃COOH CO₂ + CH₄                         (13)

Methane is poorly soluble in water and will accumulate to high levels in wetland soils and aquatic sediments. If you have ever noticed bubbles escaping from a wetland soil or a lake sediment, you have seen an important way in which dissolved CH₄ reaches the atmosphere. Methane also diffuses to the atmosphere through water and even the hollow stems of wetland plants. By either diffusion pathway, the CH₄ will even eventually encounter an aerobic zone where it is consumed by methane-oxidizing bacteria (i.e., methanotrophic bacteria):

CH₄ + O₂ CO₂ + 2H₂O                     (14)

This reaction occurs near the sedimentwater interface in submerged sediments, near the surface of the water table in exposed sediments and soils, and at the root surface of wetland plants. Because the stems of most wetland plants are somewhat porous, atmospheric O₂ diffuses through the plant and into the soil. Microbial methane oxidation consumes 5090% of gross CH₄ production. Methanotrophic bacteria in upland soils consume CH₄ directly from the atmosphere. The portion of CH₄ production that escapes microbial oxidation reacts with oxidized gases in the atmosphere to form CO₂. Oxidation of CH₄ to CO₂ completes the methane cycle (see Figure 7 for an example in natural wetlands).

Figure 7. A simplified schematic of methane cycling in a wetland ecosystem. The formula CH₂O is a generalized expression for organic carbon that originates in plants. The CH₂O is processed by communities of fungi and bacteria, eventually becoming available to methanogenic bacteria as either acetate (equation 12) or H₂ (equation 13). Note that the O₂ used by methane-oxidizing bacteria may diffuse from the atmosphere directly into the sediment surface (as shown) or indirectly through plant stems and out the roots.

Nitrous oxide is increasing in the troposphere at an annual rate of about 0.3%. It is a powerful greenhouse gas (Table 3) and participates in stratospheric ozone destruction. The only described sink for N₂O is destruction in the stratosphere, which explains its long lifetime in the atmosphere, 120 years.

Most N₂O emissions are a by-product of nitrification and denitrification by microorganisms (Figure 5). Any process that increases these processes will also stimulate N₂O emissions. Nitrification is increased by fertilization with NH₄⁺ or urea, and by clearing of forests. High rates of deforestation in the tropics may be contributing to rising atmospheric concentrations of N₂O.

Denitrification is perhaps a larger source of N₂O than nitrification. Many denitrifying bacteria, such as the well-studied members of the genus Pseudomonas, can respire by using O₂ for a terminal electron acceptor when it is available, but switch to using NO₃^- when O₂ concentrations are low. The process is most rapid in soils and sediments that are anoxic with an abundance of the required substratesNO₃ and organic carbon. Such conditions occur in many wetlands, highly fertile lakes, estuaries, and the coastal oceans. However, the process occurs even in small anoxic zones within largely aerobic environments such as upland soils and the open ocean.

Although rates of N₂O production in wetlands are high, a relatively small global area limits their contribution to 30% of total natural N₂O sources globally. The reverse is true of upland forests, grasslands, and oceans that account for about 70% of natural sources. Anthropogenic activities such as cultivation, fires, fossil fuel burning, and cattle raising have increased sources by about 50%.

7. Human Influence on Global Biogeochemical Cycles   

Human activity has touched every corner of the biosphere and threatens to reduce the life support capacity of Earth. Many of these impacts are mediated through changes in global element cycles. In many cases, the impact is due to an increase in exchange between element reservoirs (Table 4). For example, CO₂ concentrations in the atmosphere increased about 1% per year during the twentieth century as carbon was extracted from long-term storage pools (forests and petroleum reservoirs) and mobilized into the atmosphere following combustion. Carbon dioxide levels in the atmosphere continue to rise because of human population growth and increased per capita demand for energy, food, and fiber. The natural rate of nitrogen fixation has nearly doubled because of fertilizer production, emissions of nitrogen-bearing gases such as NO (nitric oxide) from fossil fuel combustion, and widespread cultivation of N₂-fixing plants such as soybean. Mining has artificially increased rock weathering rates of phosphorus and numerous trace metals. Finally, there are whole classes of compounds that never before existed and now constitute novel subcycles within the carbon cycle. Humans have accelerated key cogs in virtually every element cycle, and these changes are propagating through the biosphere and climate system.

Table 4. Examples of human intervention in the global biogeochemical cycles of carbon, nitrogen, phosphorus, sulfur, water, and sediments
Data are for the mid-1900s.

Global climate change is perhaps the most dramatic example of the level to which human impacts on the planet have grown. It is a fact that human activity during the twentieth century increased the concentrations of several important greenhouse gases. Mean global temperature increased in the same time period, though the increase did not closely track greenhouse gas concentrations. Many characteristics of the observed global warming implicate human activity as a major contributing cause.

Is it possible to prevent further global warming in the twenty-first century by reducing greenhouse gas emissions? In principle the answer is affirmative, but in practice it will require difficult economic, social, and political decisions. Knowledge of global biogeochemical cycles can help inform the decision-making process. For example, it will take decades for reductions in CO₂ emissions to influence the atmospheric concentration of CO₂. This is because the average residence time of carbon in terrestrial systems is about a century, due to the long turnover times of some pools such as wood. Thus, reductions in CO₂ emissions must begin well before actual reductions in greenhouse gas are realized.

The transition to a warmer climate will disrupt ecosystems because temperature and precipitation determine many key characteristics such as the dominant vegetation and soil development. However, we should be equally concerned about the disruptions to climate that will be caused by changes in ecosystems. We now appreciate that ecosystems control many aspects of climate at regional and global scales. Life influences climate by regulating (a) greenhouse gas concentrations, (b) coupled exchange of water and energy with the atmosphere, (c) albedo, and (d) surface roughness affecting the momentum of weather fronts moving across land. Through these mechanisms, biota can either amplify or dampen human-induced changes in radiative forcing (Figure 8). For example, an increase in wood production caused by elevated CO₂ would be a negative feedback on radiative forcing, while an increase in CO₂ emissions from forest fires would be a positive feedback. We are only beginning to understand the full extent of biotic feedbacks on climate, but they may help explain why Earth has remained habitable to life for at least 3.6 10⁹ years.

Figure 8. Anthropogenic climate forcing

The degree to which life controls the environment of the planet, and the mechanisms by which this occurs, is an active area of scientific inquiry. The Gaia hypothesis, originally proposed by James Lovelock, has stimulated speculation that life has long regulated climate within a narrow life-support range, and that this control operates through natural selection. Whether or not the Gaia hypothesis survives the test of time, it is clear that life has largely co-opted all the relevant aspects of element cycles on Earths surface, and that these cycles distinguish Earth from our lifeless neighbors in the solar system.

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Glossary   

Aerosols: Micrometer-size solid or liquid particles suspended in the air.

Albedo: The fraction of solar radiation striking a surface that is reflected rather than absorbed.

Asthenosphere: The rock layer below the lithosphere where convection currents occur that drive the movement of lithospheric plates.

Autotrophs: Organisms that obtain their energy from the sun or an inorganic chemical source. Compare with heterotroph.

Carbon quality: Used to describe the susceptibility of organic matter as a substrate for microbial decomposition and metabolism. High quality carbon compounds are easily decomposed.

Eutrophication: A process whereby a body of water becomes nutrient rich and characterized by high productivity, high microbial metabolism, and low dissolved oxygen levels.

Gross primary production (GPP): The total amount of energy gained by autotrophic organisms in the process of photosynthesis or chemosynthesis.

Global warming potential (GWP): The potential for a gas to cause radiative forcing of climate. Although somewhat controversial, the GWP of a well-mixed gas can be regarded as a first-order indicator of the potential global mean temperature change due to that gas relative to CO₂.

Heterotrophs: Organisms that obtain energy in an organic form from other organisms (compare autotroph).

Igneous rock: Rock formed by the cooling and crystallization of magma. The crystal structure and chemical composition is similar to that when it formed.

Lithosphere: The uppermost layer of the planet. It is rigid, ~100 km thick, and carries several rock plates that move about Earth forming ocean basins and continents.

Mass flow: The net movement of a volume of gas, liquid, or solid.

Metamorphic rock: Rock derived from preexisting rocks, but differing in physical, chemical, and mineralogical properties due to extreme heat and pressure originating within the earth. The preexisting material can be igneous, sedimentary, or another form of metamorphic rock.

Mineralization: The conversion of elements bound in organic compounds to an inorganic form. The term is also used to describe the formation of rocks in the geological sciences.

Net primary production (NPP): The net amount of gross primary production that remains after a portion is consumed in autotrophic respiration.

Net ecosystem production (NEP): The net amount of gross primary production in an ecosystem that remains after portions have been consumed by all the autotrophs and heterotrophs in the system.

Nitrification: Biological oxidation of ammonium (NH₄⁺) to nitrite (NO₂) or nitrate (NO₃).

Nitrogen fixation: Conversion of dinitrogen (N₂) to ammonia and ultimately organically bound nitrogen, a form that can be used by organisms.

Photosynthetically active radiation: Light wavelengths that can be used in photosynthesis (those between 400 and 700 nm).

Radiative forcing: A change in the average net radiation (solar or terrestrial) exchanged across the top of the troposphere.

Recalcitrant: Compounds or structures that decompose slowly.

Reductionoxidation: The simultaneous loss of electrons by one element (oxidation) and gain of electrons by another element (reduction).

Respiration: Energy lost by combustion of chemical energy to support metabolism.

Sedimentary rock: Rocks that develop from sediment deposits as they become more or less consolidated due to burial and aging.

Siderophores: Compounds secreted by microorganisms and some plants that form a highly stable compound with iron. They facilitate the uptake of this important trace element.

Stratosphere: The region of the atmosphere above the tropopause to ~50 km above earth. This is where the ozone layer occurs.

Subduction: One lithospheric plate is either overridden by another or forced to descend into the asthenosphere following a collision between plates.

Troposphere: The lowest region of the atmosphere (813 km high) where most air pollution and weather phenomena occur.

Upwelling zone: The upward movement of ocean water from depths of about 50150 m to the ocean surface.

Weathering: The various physical and chemical changes that occur in rocks when they are exposed to organisms and the atmosphere.

Butcher S.S., Charlson R.J., Orians G.H., and Wolfe G.V., ed. (1992). Global Biogeochemical Cycles, 379 pp. San Diego, California: Academic Press. [This is an introduction to biogeochemical cycles with an emphasis on geochemistry.]

Fenchel T., King G.M., and Blackburn T.H. (1998). Bacterial Biogeochemistry: The Ecophysiology of Mineral Cycling, 307 pp. San Diego: Academic Press. [This focuses on the important links between biogeochemical phenomena, microbial metabolism and ecology.]

Mackenzie F.T. (1998). Our Changing Planet: An Introduction to Earth System Science and Global Environmental Change, 486 pp. Upper Saddle River, New Jersey, USA: Prentice Hall. [This is a balanced introduction to biogeochemical cycles. It includes study questions but few citations.]

Schlesinger W.H. (1997). Biogeochemistry: An Analysis of Global Change, 443 pp. San Diego: Academic Press. [This is an introduction to biogeochemical cycles with an especially lucid treatment of biological processes. The text has a wealth of literature citations.]

J. Patrick Megonigal is a Senior Research Scientist at the Smithsonian Environmental Research Center in Edgewater, Maryland, USA. His research is focused on the carbon cycle as a means of understanding the responses of ecosystems to stressors such as flooding and climate change. Element cycling interactions between plants and microorganisms, particularly in wetland ecosystems, is a common theme in his field and laboratory research projects. Dr. Megonigal is a member of the Ecological Society of America, the American Geophysical Union, and the Society of Wetland Scientists. He holds an MS in biology from Old Dominion University, Virginia and a PhD in biogeochemistry from Duke University, North Carolina.