Global Biogeochemical Cycling
Summary and Keywords
The tendency to represent natural processes as cycles—from Latin cyclus and Greek κυκλος—is undoubtedly rooted in the human observations of repeating or periodic phenomena. The oldest notions of the water cycle, as water cycling between the Earth, air, and back to earth, are mentioned in the Old Testament and by Greek philosophers, from the 900s to 300s bce. The life of plants, deriving their constituents from the soil and air, and returning them thereto, is a classic example of a cycling or recycling process. For chemical elements, the concept of their cycling developed gradually since 1875 to about 1950, as the knowledge of the parts of the Earth—its compartments or reservoirs—progressed and the flow of material between them became better understood.
The main “bioessential” chemical elements are carbon (C), nitrogen (N), phosphorus (P), oxygen (O), and hydrogen (H). These are represented in the mean composition of aquatic photosynthesizing organisms as the atomic abundance ratio C:N:P = 106:16:1 or as (CH2O)106(NH3)16(H3PO4). In land plants, estimates of mean composition vary from C:N:P = 510:4:1 to 2057:17:1. On land, the photosynthesizing organisms are much more efficient than in water by being able to incorporate more carbon atoms for each atom of phosphorus. The bioessential elements are coupled by the living organisms in the exogenic cycle, the processes at and near the Earth’s surface, and in the endogenic cycle of the processes that include subduction into the Earth’s interior and return to the surface. The main reservoirs of the bioessential elements are very different: although oxygen is the most abundant element in the Earth’s crust, most of it is locked in silicate minerals as SiO2, and the forms available to biogeochemical cycling are oxygen in water and, as a product of photosynthesis, as gas O2 in the atmosphere. Carbon is in the atmospheric reservoir of CO2 gas and dissolved in ocean and fresh waters. The main nitrogen reservoir is the molecular N2 in the atmosphere and oxidized and reduced nitrogen compounds in waters. Phosphorus occurs in the oxidized form of the phosphate-ion in crustal minerals, from where it is leached into the water.
The natural cycle of the bioessential elements has been greatly perturbed since the late 1700s by human industrial and agricultural activities, the period known as the Anthropocene epoch. The increase in CO2, CH4 and NOx emissions to the atmosphere from fossil-fuel burning and land-use changes has rapidly and strongly modified the chemical composition of the atmosphere. This change has affected the balance of solar radiation absorbed by the atmosphere—generally known as “climate change”—and the acidity of surface-ocean waters, referred to as “ocean acidification.” CO2 in water is a weak acid that dissolves carbonate minerals, biogenically and inorganically formed in the ocean, and it thus modifies the chemical composition of ocean water. Overall, a major anthropogenic perturbation of the biogeochemical cycles has been the faster increase in atmospheric concentration of CO2 than its removal from the atmosphere by plants, dissolution in the ocean, and uptake in mineral weathering.
Global biogeochemical cycles of the elements and compounds are a subject of considerable scientific and public interest today. In any discussion of environmental change at any temporal or spatial scale, one must consider the substances of interest that are transported in cycles about the earth’s surface—through its air, water, soil, rocks, ice, and living and dead organic matter. The term “biogeochemical” reflects the fact that biological, geological, chemical, and physical processes play important roles and interact with each other in the global material cycles that are mediated to a large degree by photosynthetic primary production and respiration or mineralization of organic matter. Thinking about these global biogeochemical cycles and their role in environmental change requires crossing over the usual discipline boundaries between biology, ecology, oceanography, meteorology, chemistry, and geology. Because of the impact of human activities on the global cycles, and consequently, for example, the climate, the subject also involves the effects and consequences of natural and human-induced change for ecosystems, humans, and human infrastructures. This can lead the discussion into the fields of sociology, economics, and political science.
In a short article such as this, it is impossible to cover the essentials of all the global biogeochemical cycles that naturally exist for the chemical elements and compounds. To begin with, the ecosphere, or exogenic and shallow endogenic system—Earth’s outer shell—is highly interactive with matter and energy flowing between and within individual ecosystems in interconnected element and compound biogeochemical cycles. Most elements and compounds on Earth are involved in these cycles, and innumerable processes cycle elements throughout the ecosphere. As an example, consider the group of microscopic single-celled organisms called the prokaryotes: the bacteria and blue-green algae. The processes that these organisms are involved with include (a) the capture of carbon dioxide from the atmosphere and its conversion to organic matter (fixation of CO2), (b) the release of CO2 back to the atmosphere through respiration and decay, (c) fermentation of sugar, (d) methane production and oxidation, (e) sulfur reduction and oxidation, (f) nitrogen fixation, (g) nitrification, and (h) denitrification. This list is given only as an example of the diversity of processes that affect the transport of materials about the Earth’s surface and drive the biogeochemical cycles of elements and compounds. For various reasons, including those involving nutrient cycling and limitation and climate control, it is essential that the integrity of all biogeochemical cycles on Earth be preserved. As illustrations of the processes, reservoirs, and fluxes involved in the global biogeochemical cycles of elements and compounds, we have chosen to look at the cycles of several elements termed “bioessential”: carbon, nitrogen, and phosphorus. The carbon cycle discussion is the most extensive since it is the foundation of the cycles of the other bioessential elements and carbon in the form of CO2 is a major variable involved in the human-induced environmental problems of climatic change and ocean acidification.
Some Brief Historical Comments on Global Biogeochemical Cycles
An early treatment of geochemical cycles appeared in 1875, when several chapters on the cycles of chemical elements were included in a book on Earth’s history by Friedrich Mohr, a professor at the University of Bonn, short chapters on the silicon and carbon cycles among them. In 1893 Professor Arvid Hogbom, a colleague of Svante Arrhenius, presented a lecture at the Physical Society drawing the conclusion that the chief source of atmospheric CO2 is in the release of the gas during the natural breakdown of limestone. He further theorized concerning the long-term geologic cycle of carbon by addressing the questions of how much of the CO2 is retained in the atmosphere, how much is absorbed by the ocean, and how changes in the concentration of this gas in the atmosphere might affect climate. By the 1920s, the cycles of the chemical elements that are involved in biological processes, the bioessential elements of carbon, nitrogen, and phosphorus, and are also transported between soil, crustal rocks, atmosphere, land and ocean waters, and the Earth’s interior, were well recognized. Alfred Lotka’s book, Elements of Physical Biology, published in 1925, has chapters on the cycles of carbon dioxide, nitrogen, and phosphorus that present a modern treatment of what we call today the biogeochemical cycles. Furthermore, he wrote that his ideas of the nutrient element cycles and mathematical treatment of biogeochemical problems had been developed as far back as 1902 and in his publications starting in 1907. Vladimir Ivanovich Vernadsky was the principal architect of our contemporary vision of the biosphere and noosphere. The noosphere is a new evolutionary stage of the biosphere, “where man, for the first time in the history of the earth, knew and embraced the whole biosphere, completed the geographic map of the planet earth, and colonized its whole surface. Mankind became a single totality in the life on earth. The noosphere is the last of many stages in the evolution of the biosphere in geological history” (Vernadsky, 1945). Vernadsky’s essays on geochemistry and the biosphere fortunately are now available in English translations (e.g., Vernadsky, 2007) so that the Western world can gain a better appreciation for the extent and creativeness of his work.
By 1950, the geochemical cycles of elements in the Earth’s interior and on its surface certainly were textbook material, appearing in Rankama and Sahama (1950). This early, if not first, systematic textbook treatment of the geochemical cycles presented diagrams of the geochemical reservoirs as boxes and fluxes between them, and tabulations of the elemental concentrations or masses in some of the individual reservoirs. Subsequent decades produced the knowledge we have today of the chemical speciation of the elements in the different compartments of the Earth, their abundances, and mechanisms responsible for their flows. Whereas the earlier models of the global biogeochemical cycles of individual elements were static, describing the cycles without their evolution in time, developments in the mathematical treatment of time-dependent multireservoir systems found their application in the analysis of biogeochemical cycles (e.g., Odum & Odum, 1959; Lerman, Mackenzie, & Garrels, 1975). Since then, there has been a great proliferation of cycle models and, in particular, of carbon cycle models at very different spatial and temporal scales, aimed at interpretation of cycle evolution in the past and its projection into the future for the world as a whole, as well as for such global reservoirs as the atmosphere, land, coastal oceanic zone, and the open ocean.
Considerable attention in the past few decades became focused on the global sedimentary cycle and the cycling of salts in the ocean, as a result of Kelvin’s (William Thomson, later Lord Kelvin) estimates of the age of the Earth between 24 and 94 mega-annum (Ma), made between 1864 and 1899 (Burchfield, 1975), and the estimates of the age of the ocean from the rate of accumulation of sodium brought in by rivers, as was done, for example, by Joly (1899) whose age of the ocean he calculated as about 90 Ma. Gregor et al. (1988) summarized and discussed in detail the geological arguments in the second half of the 1800s and the early 1900s for the recycling of oceanic sediments after their deposition and for the existing sinks of dissolved salts in ocean water, such as their removal by adsorption on clays, entrapment in sediment pore water, and formation of evaporites that were contrary to the idea of the ocean continuously filling up with dissolved salts, as thought by Joly. Garrels and Mackenzie (1971) presented the concepts of the sedimentary cycling of materials, which had lain dormant for some years, in book form, and in 1972 these two authors developed a quantitative model of the complete sedimentary rock cycle. Later work (Veizer, 1988) showed that the recycling rates of the sedimentary lithosphere and the various rock types within it are mainly a function of the recycling rates of the tectonic realms, such as active margin basins, oceanic crust, and continental basement, in which the sediments were accumulated. The sedimentary cycle of the elements is discussed in more detail in “Global Sedimentary Cycle of the Elements.”
Construction of Global Biogeochemical Cycle Box Models
The most common methodology used in studying the global movements of elements and compounds is by mathematical modeling of biogeochemical cycles at the Earth’s surface. Modeling also allows scientists to estimate the effects of human activities on natural biogeochemical cycles. Biogeochemical cycling models generally include processes that move materials and their rates of transfer among a limited number of well-studied spheres of the Earth. Biogeochemical cycles, however, have certain properties that are inherently difficult to describe and model. These include (a) irreversibility, that is, the system does not return to its exact previous state if it goes through a disturbance; (b) transitional phenomena, that is, the system tends to switch from one state to another and yet others, and perhaps back again, rather than simply moving from “before” to “after”; (c) evolution, in which the system progressively changes in a particular direction; and (d) processes that either enhance the original perturbation to the system (positive feedback) or relieve the perturbation (negative feedback).
Biogeochemical cycles of elements and compounds are often portrayed in the form of box models. In developing a biogeochemical cycling model, it is necessary to separate first the system of interest from its natural surroundings. In general, the boundaries of a natural system are defined by the scale of the phenomena of interest and by previous knowledge of possible interactions between the system and its surroundings. Global biogeochemical cycling models involve consideration of phenomena on a worldwide scale. The Earth is divided into a number of physically well-defined spheres referred to as “boxes” or “reservoirs.” The term “box model” is commonly applied to models of this type.
In Figure 1 of the description of a global biogeochemical box model, three reservoirs of an hypothetical substance A are shown with transport paths of A between reservoirs. The ocean reservoir is subdivided into the compartments of water, biota, and sediment. The land reservoir includes the compartments of plants, animals, and soil. The masses of the substance A in these compartments are known. However, although transport paths among the compartments are suspected, no values for the fluxes are known for these flow paths. The flows between the reservoirs involve the masses transported. The rates of transport, fluxes, are generally measured in units of mass per unit of time. The fluxes may be related to physical, chemical, geological, or biological processes. The quantitative evaluation of their contribution is probably the most critical point of model building. It necessitates both a good knowledge of all the interactions possible between the reservoirs and the dynamics of the processes involved. The complexity of the task increases rapidly if subsystems are considered. In this case, it is necessary to understand the physical processes of movement, such as mixing by advection and diffusion, as well as biological processes, within the reservoirs. Herein lies a major problem in the modeling of global cycles of the elements. The root cause is that, in many cases, basic data and an understanding of the processes operating in the system of interest are lacking. Many material cycles have innumerable processes associated with fluxes. Some are better known than others, and some are quantitatively unknown. Our knowledge of the global biogeochemical cycles, particularly their dynamics and potential for feedback during perturbations, is still rudimentary for many elements but is expanding rapidly.
The transport of substance A involves (a) gaseous transport from the land and sea surface to the atmosphere (evasion, F12 and F32) and return in rainfall (F21) and precipitation (F23); (b) river transport of dissolved and solid materials (F13); and (c) return of the substance to the land via uplift of sedimentary rocks (F31). Fluxes of substance A are known for these transport paths. Any transformations or chemical reactions, such as oxidation or reduction and solution or precipitation, are not represented in the diagram of the cycle. The quotient dM/dt simply represents the change in mass of a reservoir per unit of time, the rate of change of mass. In a steady-state system, it is equal to zero. In a non-steady-state system, it differs from zero. Knowledge of the fundamental processes and of the driving forces behind the flows may not be sufficient to enable development of a quantitative relationship between them and the material fluxes. Thus, the fluxes of materials (denoted Fij in the figure for the flux from box i to box j) are commonly measured and related to the conditions in the system according to some chosen model. The two simplest flux models are of zeroth- and first-order fluxes. The zeroth-order flux is a constant:
A first-order flux is one that is proportional to the reservoir mass:
where Mi is the mass of a substance in a reservoir i, and kij is a rate parameter for the flux going from reservoir i to j. In general, kij may vary with reservoir size and time, and may be a function of environmental conditions within a system. However, in the preceding equation, it is treated as a constant. In a steady-state (unchanging) system, reservoir concentrations or masses of a substance do not change with time. This requires that the input and output fluxes for every reservoir are equal. If one of the fluxes changes, the steady state of the system becomes perturbed. The system is no longer at steady state, and the system is described as transient or non-steady-state. Such a condition can result in changes in all the reservoir masses.
The residence time (λ) of a substance in a reservoir provides some clue as to the reactivity of a substance within a reservoir. It is defined as the ratio of the mass of the reservoir to the sum of either input or output fluxes of the substance at steady state. λ is equal to 1 divided by the rate constant:
A perturbation of a biogeochemical cycle caused by a change in an input flux will result in a change in reservoir mass. For a fixed change in input, up or down, the reservoir mass will come to within 5% of a new steady-state value after three residence times have elapsed. Thus, most of the change (95%) caused by a perturbation in input would be completed in a time period approximately equal to three times the residence time. Consequently, reservoirs of short residence times respond rapidly to external perturbations. In reservoirs of long residence time, perturbations require more time to work their way through.
Feedback is an important concept in biogeochemical cycles and environmental change. Feedback is a self-perpetuating mechanism or process of change and response to that change. Natural systems react to perturbations in a positive or negative fashion. In a positive feedback loop, the effects of a perturbation are amplified; in a negative feedback loop, the effects of the disturbance are diminished.
The Global Sedimentary Cycle of the Elements
In 1972 R. M. Garrels and F. T. Mackenzie wrote a paper entitled “A Quantitative Model for the Sedimentary Rock Cycle,” which has formed the basis of our modern understanding of the sedimentary cycle of the elements. This model described the steady-state cycling of 11 elements important in the formation and destruction of sedimentary rocks: Al, C, Ca, Cl, Fe, K, Mg, Na, S, Si, and Ti. The model was based on a steady state mass balance of the elements in the atmosphere, ocean, and sedimentary lithosphere that was consistent with the observed composition of the atmosphere, ocean, stream, and groundwater reservoirs, as well as the mass–age distribution of the chemical and mineralogical composition of sedimentary rocks. Conditions were incorporated in the model that allowed for rapidly recycling and more slowly recycling sedimentary reservoirs with material transfers between each due mainly to post-depositional diagenetic processes. The model was secular and cannibalistic and did not allow for inputs involving endogenic processes that act to reconstitute primary igneous crystalline phases and volatile acids, such as CO2, from sediments, organic matter, and dissolved constituents. This implied that there was little exchange of materials between the sedimentary lithosphere and the endogenic cycle involving the mantle.
The description of the sedimentary cycle of the elements has become more detailed and quantitative with time and now incorporates interactions with the mantle. Its details are difficult to show in a single illustration. However, in Figure 2 is a schematic representation of one of the subcycles of the sedimentary rock cycle, that of the calcium-magnesium-silicate-carbonate-carbon dioxide system.
This subcycle is important to the controls on paleoatmospheric CO2 during the past 545 Ma of Earth’s history (the Phanerozoic, Figure 3). As portrayed in Figure 2, calcium carbonate (CaCO3), along with some organic carbon, can be subducted to the mantle or removed to a metamorphic regime, such as deep burial in sedimentary basins like the Gulf Coast of the United States. Both of these pathways lead to decarbonation reactions that serve to return CO2 originally consumed in the weathering cycle to the atmosphere. The inorganic reaction is the classic Urey-Ebelmen reaction simply written:
The SiO2 represents original silica deposited as biogenic opal-A and transformed to quartz (SiO2) during burial and subduction.
Owing to these processes and others involving mainly the evolution of the biosphere, paleoatmospheric CO2 has varied during geologic time (Figure 3). Complementary to these atmospheric CO2 concentration variations are accompanying changes in the chemistry of the ocean (Figure 4), chemistry and mineralogy of biogenic and inorganic carbonate precipitates from seawater, and that of carbonate sediment types throughout Phanerozoic time (e.g., Mackenzie & Andersson, 2013). These temporal changes in paleoatmospheric CO2 and ocean chemistry are a result of processes and changes in the magnitudes of the fluxes associated with them mainly within the calcium-magnesium-silicate-carbonate-carbon dioxide system of the global sedimentary cycle.
Global Biogeochemical Cycle of Carbon
Late Preindustrial Time
Life on Earth is based on carbon as one of the main components of organic matter. The occurrence of the various forms of carbon in different parts of the Earth’s interior and its surface shells, and the processes that are responsible for the transfer of carbon between the different parts of the Earth form a conceptual model of the geochemical or biogeochemical cycle of carbon. The carbon cycle is usually divided into a deeper part, called the endogenic cycle, and the part that includes the surface reservoirs of the sediments, oceanic and continental waters, land and aquatic biomass, soils, and the atmosphere is referred to as the exogenic cycle. The main features of the global carbon cycle are shown in Figure 5 with the carbon inventory on Earth. The inventory values of the reservoirs represent those of late preindustrial time, with the atmospheric carbon dioxide mass equivalent to a preindustrial level of CO2 of 280 ppmv, which, as shown in Figure 3, varied significantly in the geologic past and has risen about 40% during the Industrial Age.
Most of the carbon in the exogenic cycle is inorganic carbon found in the sedimentary rocks of limestone and dolomite. The ratio of inorganic carbon to organic carbon in sediments is ~5.2. Concentrated organic deposits of coal, oil, natural gas, mainly as methane (CH4) and methane hydrate (clathrate of water ice and methane), constitute only 0.02% of all the carbon in sediments and sedimentary rocks. However, it is the burning of this carbon, mainly as an energy source for human society, and the resultant emissions of CO2 to the atmosphere that are mainly responsible for rising global atmospheric CO2 concentrations during the past ~150 years. The anthropogenic emissions of CO2 to the atmosphere have led to the problems of global warming of Earth’s surface and the acidification of surface ocean waters.
The global soil reservoir contains mainly organic carbon in the form of nutrient-rich humus consisting of decayed plant and animal matter. Inorganic carbonate constitutes about 25% of the total carbon in the soil reservoir. In contrast, carbon in ocean water is present mainly in the form of dissolved inorganic carbon (DIC), containing the chemical species of bicarbonate (HCO3−) and carbonate (CO32−) anions, and dissolved molecular CO2. Total organic carbon (TOC) in seawater present as particulate organic carbon (POC) and dissolved organic carbon (DOC) makes up only 2.6% of the total carbon in the ocean reservoir. It should be noted that the DIC carbon reservoir mass in the ocean exceeds the atmospheric carbon reservoir mass of CO2 by 47 times! Thus, small changes in the mass of the former due to changes in net primary production (NPP), for example, and subsequent changes in the air-sea exchange of CO2 can have a significant impact on the size of the atmospheric reservoir of carbon. The land biomass carbon reservoir of living plants, present mainly as forest trees with lifetimes of several decades, far exceeds the carbon mass of primary producers in the ocean, mainly the phytoplankton, by 240 times, and the latter turns over at a rate that is ~1500 times faster than the former.
Besides the portrayal of the globally important carbon reservoirs in Figure 5, the major processes moving carbon through the global biogeochemical cycle of the element are also shown in the figure. Fluxes of the element between reservoirs associated with these processes are shown in Table 1. Total global net primary production is equivalent to ~9500 × 1012 mol C/yr (113.5 Gt C/yr); gross primary production (GPP) is about double this value. Of Net Primary Production (NPP), 55.6% is on land and 44.4%, mainly as phytoplankton, is in the ocean. Another large carbon flux value in the global carbon balance involves the decay and volatilization of organic matter in soils and release to the atmosphere mainly as CO2. This flux is equivalent to 5200 × 1012 mol C/yr (62.5 Gt C/yr). On land 22 × 1012 mol C/yr (0.26 Gt C/yr) are involved with the weathering of silicate and carbonate sediments and rocks resulting in riverine transport of DIC equivalent to 32 × 1012 mol C/yr (0.38 Gt C/yr). The carbon flux transported by rivers to the ocean as total organic carbon (POC + DOC) is nearly equivalent to that entering the ocean as DIC. The land to ocean movement of carbon via river and groundwaters forms a continuum of processes termed the Land Ocean Aquatic Continuum (LOAC) that play an important role in the uptake of anthropogenic CO2 in the terrestrial realm and its fate thereafter (Regnier, Friedlingstein, Ciais, & Mackenzie, 2013).
In the late preindustrial view of the global carbon cycle and its fluxes (Figure 5 and Table 1), the ocean took up about 8000 × 1012 mol C/yr (96 Gt/yr) in organic productivity and physical dissolution of the gas in the surface waters of the ocean and released through respiration and decay of organic matter and precipitation of calcium carbonate, mainly in the biogenic shells and tests of calcifying organisms, approximately 42 × 1012 mol C/yr (0.51 Gt C/yr) more than it absorbed. Although a somewhat controversial conclusion, it appears that the oceans were a net source of CO2 to the atmosphere in preindustrial times. The land absorbed the net flux of CO2 from the ocean in terrestrial productivity and sequestration of organic carbon in soils.
On the geologically long term, ~23 × 1012 mol C/yr (0.28 Gt C/yr) are sequestered on the seafloor, 18 × 1012 mol C/yr (0.22 Gt C/yr) in accumulation of carbonate minerals and 5 × 1012 mol C/yr (0.06 Gt C/yr) in organic matter. Thus, only 20% of total CaCO3 production of the minerals aragonite (orthorhombic CaCO3), calcite (hexagonal CaCO3), and magnesian calcite (hexagonal Ca1-xMgxCO3) compositions and 0.1% of NPP in surface waters are eventually sequestered in the sediments of the seafloor. In most recent times, the accumulation of carbonate on the seafloor is given as 31 × 1012 mol C/yr (0.37 Gt C/yr) and that of organic carbon as 16.3 × 1012 mol C/yr (0.196 Gt C/yr) (cf. Andersson & Mackenzie, 2004; Smith & Mackenzie, 2015).
The large differences between production and seafloor accumulation reflect the fact that much of the organic carbon produced in surface ocean waters is microbially respired by oxygen in the water column as organic matter sinks toward the seafloor releasing CO2 to the water column. The CO2-enriched aggressive waters at mid and deep depths in the ocean dissolve the carbonate in transit to the seafloor. Smith and Mackenzie (2015) conclude that the net flux of CO2 associated with CaCO3 reactions in the global ocean is an apparent influx from the atmosphere to the ocean due to dissolution of carbonate of 7 × 1012 mol C/yr (0.08 Gt C/yr) at a time scale of 102–103 years. The authors also conclude that the CaCO3-mediated influx of CO2 is approximately offset by CO2 release from organic C oxidation in the water column. This is a remarkable balance of the carbon system in the ocean.
To complete the picture of the preindustrial global carbon cycle, the processes of volcanism, metamorphism, and hydrothermal reactions release 18 × 1012 mol C/yr (0.22 Gt C/yr) to the exogenic system, and uplift of the carbonate rocks limestone and dolomite return 33 × 1012 mol C/yr (0.40 Gt C/yr) to the Earth’s surface for weathering annually.
No planetary biogeochemical system is in an absolute balance of all inputs and outputs on a global scale. Figure 5 and Table 1 are our best estimates of the nature of the quasi-steady-state global carbon cycle balance of late preindustrial time. It is this carbon system that is being significantly perturbed by the activities of humankind. In the next section keeping in mind the discussion of this section, we consider in detail the more recent global carbon cycle as perturbed by fossil-fuel burning and land-use changes.
Modern CO2 Cycle and Climatic Change
In this section, we will discuss the modern global carbon dioxide cycle focusing on the exchange of carbon between Earth’s surface and its atmosphere. Perturbation of this exchange due to human activities has led to an enhanced greenhouse. The planet throughout most of its evolutionary history has always had a greenhouse effect first recognized by Joseph Fournier in 1824, but not called such until 1901 by Nils Gustaf Ekholm (Ekholm, 1901; Crawford, 1997; Fleming, 2001, 2005). The natural greenhouse effect is a set of processes by which atmospheric gases of H2O, CO2, CH4, and N2O, termed “greenhouse gases,” absorb long-wave Earth radiation radiated from the planetary surface and reradiate the radiation in all directions warming the near surface of the Earth. Without the radiatively active gases in the atmosphere, Earth’s lower atmosphere and surface would be approximately 32oC colder than its late preindustrial global temperature of 14oC. The intensification of the natural greenhouse effect by release and storage of some amount of the emissions of greenhouse gases in the atmosphere due to human activities, like fossil-fuel combustion, has led to an increase in the strength of the planetary greenhouse. This enhancement is known as the enhanced greenhouse effect.
Building on the contributions of Joseph Fourier, John Tyndall, and Claude Pouillet, the Nobel Prize–winning Swedish scientist Svante Arrhenius in 1896 calculated how changes in the concentration of carbon dioxide in the atmosphere are related to the surface temperature of the Earth. He concluded that “if the quantity of carbonic acid [CO2] increases in geometric progression (in the atmosphere), the augmentation of the temperature will increase in arithmetric progression” (Arrhenius, 1896). This conclusion can be formulated in the simple equation:
where ΔF is the radiative forcing in watts per square meter, α is a constant with a value between 5 and 7, C is the concentration of atmospheric CO2 in ppmv, and Co is its concentration at a baseline level. Borrowing from the carbon cycle studies of his Swedish colleague Arvid Hogbom, Arrhenius was the first person to calculate the effect of rising atmospheric CO2 concentrations due to CO2 emissions from fossil-fuel burning and other combustion processes on the temperature of the Earth’s surface. His conclusion was that a doubling of atmospheric CO2 would lead to a rise of global mean temperature of about 4.5oC, very close to the upper limit of the range in temperature estimated for a doubling of atmospheric CO2 from modern global climatic models.
The major form of carbon in the atmosphere is CO2(g), constituting more than 99% of atmospheric carbon, and the most important trace gas emitted by human activities to the atmosphere contributing to the potential of an enhanced greenhouse effect is CO2. Figure 6 is a representation of the global biogeochemical cycle of carbon dioxide emphasizing the CO2 exchanges between the Earth’s surface and atmosphere. The atmospheric CO2 concentration and its mass are values for the early part of the twenty-first century. In April 2016, the CO2 concentration in the air above the trace gas observatory on Mauna Loa, Hawaii, had reached slightly more than 407 ppmv. This is the highest concentration in human history and probably for the past 10 to 15 Ma (Tripati, Christopher, & Robert, 2009). The rate of atmospheric CO2 growth was 3.05 ppmv/yr in 2015. Atmospheric CO2 has risen ~30% over its late preindustrial level of 280 ppmv. Figure 7 shows the Mauna Loa Observatory record for atmospheric CO2 concentrations from 1958 to 2015.
A number of the natural CO2 fluxes shown in Figure 6 were discussed in “Late Preindustrial Time.” The most important additions to these fluxes are those involving human activities of fossil-fuel combustion and land-use changes. In the early 21st century, 0.64 × 1015 mol C/yr (7.7 Gt C/yr) and 0.12 × 1015 mol C/yr (1.4 Gt C/yr) were emitted to the atmosphere from the burning of coal, oil, and gas; in the process of cement manufacturing (about 2% of total industrial emissions of CO2); and in land-use changes, such as deforestation, respectively. The industrial emissions from the most recent estimations of the Global Carbon Budget project were 0.81 × 1015 mol C/yr (9.74 Gt C/yr) in 2015 and from land-use changes 0.092 × 1015 mol C/yr (1.1 Gt C/yr) in 2014 (Le Quéré et al., 2015). The average for the period of 2005–2014 fate of the total anthropogenic emissions of 0.83 × 1015mol C/yr (9.9 Gt C/yr) was that 44% remained in the atmosphere, and 30% and 26% were taken up in the terrestrial realm and ocean, respectively; that is, the anthropogenic carbon missing from the atmosphere, shown in Figure 6 as “missing,” was partitioned into the land and ocean reservoirs. The uptake of anthropogenic carbon in both reservoirs of land and ocean represents only a very small fraction of the total exchange of carbon between the Earth’s surface and the atmosphere.
An older but still notorious and talked about problem with atmospheric CO2 today is the difficulty in balancing the known sinks with the anthropogenic sources from fossil-fuel combustion and land-use activities, such as biomass burning. The problem is illustrated in Figure 6. We know that in the early twenty-first century, the anthropogenic flux of CO2 from fossil-fuel burning plus cement manufacturing and land-use changes, such as deforestation, was equivalent to 0.81 × 1015 mol C/yr (9.7 Gt C/yr). Anthropogenic CO2 has been accumulating in the atmosphere at a rate equivalent to 0.36 × 1015 mol C (4.3 Gt C) annually. We also know that the amount of carbon taken up by the surface ocean has been about 0.22 × 1015 mol C/yr (2.6 Gt C/yr). Subtracting the summation of the known atmospheric and oceanic accumulation of carbon from the known anthropogenic source flux, we see that the fate of roughly 0.25 × 1015 mol C/yr (3 Gt C/yr) of human-produced carbon is presumably unaccounted for. These numbers suggest that there is another important sink (sometimes called the “missing sink”) of anthropogenic CO2 besides the atmosphere and ocean.
The resolution of this apparent imbalance in the global carbon cycle lies with the land. While the deforestation of tropical rain forest ecosystems occurs at rates of about 1% of the world’s forested area per year, and other land-use activities are sources of carbon to the atmosphere, higher-latitude terrestrial ecosystems may be sinks of carbon. Such a net sink implies that some processes of carbon storage on land have changed in recent decades. Some or all of the following changes may have occurred: (a) increased atmospheric levels of CO2 act as a fertilizer and stimulate productivity in plants. This leads to storage of carbon in biomass or in soil organic carbon; (b) plant productivity is stimulated by increased NO3 and NH4 from fertilizers used in farming and from deposition of anthropogenic N from the atmosphere. Carbon once more is stored in biomass or in soil; and (c) vegetation regrowth in previously disturbed ecosystems, or growth in undisturbed ones. The first process on this list, carbon dioxide fertilization, is a potentially strong negative feedback on global warming. We know from experiments on plants and small vegetated ecosystems that almost all agricultural crops and some perennial plants, when subjected to increased CO2levels, will increase their rates of photosynthesis and growth. If this enhancement should also occur in large ecosystems, like forests, CO2would be withdrawn from the atmosphere and stored in plant organic matter. Humans have put excess phosphorus, nitrate, and ammonia nutrients into the environment by applying fertilizers to the land surface, burning fossil fuels and biomass, and discharging sewage containing nitrogen and phosphorus. These nutrients can stimulate increased plant growth in the soil and aquatic environments. This enhanced eutrophication of both land and marine environments is a negative feedback on accumulation of carbon in the atmosphere from anthropogenic sources and hence is a negative feedback on warming of the Earth. Total land and ocean eutrophication may amount to a billion tons of carbon per year (e.g., Woodwell et al., 1998).
The first problem of anthropogenic CO2 emissions involves the climate and is often termed global warming. From 1765 to the early twenty-first century, CO2 accounted for about 60% of the human-induced global warming of the Earth due to the accumulation of all greenhouse gases in the atmosphere from human activities (Figure 6). A doubling of CO2 in the atmosphere over its late preindustrial of 280 ppmv could lead to an about 3oC increase in global mean surface temperature. Depending on how sensitive the climate is to rising concentrations of anthropogenic greenhouse gases and other factors, their atmospheric accumulation could eventually lead to an increase in global mean temperature of less than 1oC to about 5oC in the year 2100 (Figure 8; Intergovernmental Panel on Climate Change, 2013). Future projections of CO2 and other greenhouse gas concentrations, and hence temperature, depend to a significant extent on the emissions scenarios used in global carbon and climatic models to calculate future gas concentrations and temperature (Figure 8). To date the global mean temperature has risen since 1850 about 0.8oC, arguably due to the enhanced greenhouse effect powered by the accumulation of greenhouse gases in the atmosphere arising from human activities. In 2013 approximately 185 countries represented by their delegations and leaders met for the twenty-first yearly session of the Conference of the Parties to the 1992 United Nations Framework on Climate Change and achieved universal agreement on climate meant to stem human-induced greenhouse gas emissions and global temperature rise. The global biogeochemical cycle of CO2and its changes during recent time constituted important background for the discussion at this conference.
The uptake or absorption of anthropogenic CO2 by the ocean (Figure 6) has led to the second problem of anthropogenic CO2 emissions, that of ocean acidification (OA). As CO2 is added to seawater, the CO3 2− ion in seawater is titrated by the added CO2 and is converted to bicarbonate ion. The overall reaction is:
and leads to a decrease in the pH of the surface waters of the ocean (increase in acidity). The concentration of carbonate ions is expected to decline by 50% during this century due to increased atmospheric carbon dioxide levels. The acidity of surface ocean waters has already increased ~30% due to the absorption of the gas CO2 generated by human activities in the ocean (Figure 9, e.g., Doney, 2006; Mackenzie & Andersson, 2013). Depending on the extent of future CO2 emissions and other factors, the Intergovernmental Panel on Climate Change (2013) predicts that ocean acidity could increase by 150% by 2100. A 150% increase in ocean acidity, and an even lessor increase, could have very serious consequences for marine ecosystems, particularly their calcifying organisms that build skeletons and tests of a variety of carbonate mineralogies. These include, for example, the pelagic, microscopic algae Coccolithophoridae, which form individual plates around them of CaCO3 (the mineral rhombohedral calcite) termed coccoliths, and the pelagic free-swimming Gastropoda pteropods (from the Greek for wing-foot), which build a skeleton of CaCO3 (the mineral orthorhombic aragonite). In the shallow-water benthic realm, organisms of note are the reef-forming corals (skeleton of aragonite) and coralline algae (skeleton of magnesian calcite, calcite with Mg substituted in its lattice). These calcifying organisms have been shown to be particularly vulnerable in terms of their calcification rate to small decreasing changes in the pH of seawater. As the pH of seawater decreases, it is anticipated from most experimental and field observations that as the carbonate ion concentration of seawater falls, the rate of calcification of marine calcifiers will decrease (e. g., Gattuso, Frankignoulle, Bourge, Romaine, & Buddemeier, 1998; Kleypas, McManus, & Menez, 1999; Riebesell, 2004; Jokiel et al., 2008; Kuffner et al., 2008; Gattuso & Hansson, 2011). However the findings with respect to the Coccolithophorid Emiliania huxleyi are somewhat inconclusive, although it appears that this pelagic calcifier can adapt to OA (Lohbeck et al., 2012, 2014).
Marine calcifiers face a second challenge of OA in addition to stress on their ability to calcify. In environments that are too acidic, the shells of calcifiers may dissolve. Coralline algae and other calcifiers constructing skeletons of magnesian calcite compositions are particularly susceptible to dissolution because of the greater solubility of these carbonate phases. It has already been demonstrated in the field that the particles produced from the disintegration of biogenic magnesian calcite organisms will dissolve in the water column and/or shallowly buried sediments if the pH is sufficiently low enough (the pCO2 is high) (Andersson, Bates, & Mackenzie, 2007; Drupp, De Carlo, & Mackenzie, 2016). Morse, Andersson, and Mackenzie (2006) referred to the biogenic magnesian calcite mineral compositions as the “canary in the cage” of ocean acidification. In addition, laboratory studies show that biogenic carbonate minerals subjected to higher than ambient CO2 concentrations will dissolve (Walter & Morse, 1985; Pickett & Andersson, 2015). Thus, it would be anticipated that under rising atmospheric CO2 conditions of the future and hence greater acidity of ocean waters that marine carbonate substrates would be more vulnerable to dissolution. Coral reef calcifiers are particularly at risk and these ecosystems have received a great deal of attention in terms of OA and rising temperatures. The latter has already produced increased incidences of coral bleaching and die off in the Great Barrier Reef of Australia, Hawaii, and Florida of the United States, and islands of the Caribbean.
Figure 10 shows one model output of changes in the carbonate production rates versus carbonate dissolution rates for shoal water carbonate ecosystems, mainly reef systems, under a BAU IS92a (Intergovernmental Panel on Climate Change, 2007, 2013) emissions scenario and consequent global temperatures and OA changes into the future. This model simulation predicts that CaCO3 dissolution will exceed CaCO3 production in about the year 2150 and that reefs will no longer be net accreting systems. The timing of this model simulation is highly uncertain, but the trend of is robust.
It should be pointed out that deep, cold, CO2-charged ocean waters are naturally too acidic for marine calcifiers to calcify and survive, and these waters and the colder surface waters of the higher latitudes have a lower carbonate saturation state (less CO32− ion) than that of the surface waters of the tropics and subtropics. With ocean acidification, the carbonate saturation horizon, which demarcates the horizon at which saturated waters (more CO3 2−ion) above are separated from undersaturated (less CO3 2− ion) waters at greater depth, is expected to shift closer to the surface. This has already been observed for waters of the eastern Pacific. In addition, waters charged with CO2 in part derived from the absorption of anthropogenic CO2 in the ocean are upwelling onto the shoal water borderland of central Canada to northern Mexico (Feely, Sabine, Hernandez-Ayon, Ianson, & Hales, 2008). These waters are undersaturated with the mineral aragonite, and the uptake of anthropogenic CO2 has already increased the areal extent of these aragonite corrosive waters in the affected area. The acidic waters may account for the lack of shells, tests, and particles of biogenic magnesian calcite compositions in the sediments of the southern California borderland, as observed by Smith (1972). Certain low carbonate saturation state surface waters in the Southern and Arctic oceans may become undersaturated with respect to aragonite near the end of this century and make it very difficult for shelled organisms of aragonite, like the pelagic pteropods, and benthic biogenic aragonitic and magnesian calcite composition organisms to calcify (Andersson, Mackenzie, & Bates, 2008).
Having looked at portions of the global biogeochemical of carbon cycle in some detail, it is now time to investigate the coupling of this cycle to the driver nutrient cycles of nitrogen (N) and phosphorus (P). Much of the discussion will center about a conceptual model of the coupled and interactive cycles of C, N, and P termed TOTEM (Terrestrial-Ocean-aTmosphere Ecosystem Model). The coupling is mainly achieved through the Redfield ratios of the terrestrial phytomass, organic matter in soils and marine sediments, and marine phytoplankton.
The Coupled Global Biogeochemical Cycles of Carbon, Nitrogen, and Phosphorus
This section continues the discussion of the global biogeochemical cycles of the elements with a look at the cycles of two elements tied to carbon with accompanying Redfield ratios through photosynthesis and respiration, those of nitrogen and phosphorus. These elements are considered “drivers” of the carbon cycle, and are often referred to as limiting or controlling nutrients. We first make some general comments about the elements nitrogen and phosphorus and then investigate how C, N, and P move throughout the ecosphere or outer shell of the Earth.
Brief Commentary on Nitrogen and Phosphorus Biogeochemical Dynamics
Nitrogen is one of the essential nutrients for organisms, and is often regarded as a limiting nutrient for biological productivity on land and in the ocean. What makes nitrogen distinct from other nutrients is (a) that the primary and largest reservoir is the atmosphere where diatomic nitrogen (N2) comprises 78% by volume, and (b) that N2 is inert in most biogeochemical processes. The large mass of the atmospheric N2 reservoir and its long residence time make the changes in its size difficult to measure. The apparent small variation of atmospheric N2O (a greenhouse gas) concentrations from 220 ppbv to 285 ppbv in the ice core records is good evidence that the nitrogen cycle varied somewhat in preindustrial time, along with variations in the cycles of CO2 and CH4 (Petit et al., 1999). However, after industrialization, the burning of fossil fuels, cultivation of legumes, and the use of organic (e.g., guano) and mineral (NaNO3, NH4NO3) nitrogen for fertilization have been major sources of reactive nitrogen input into the landscape. These anthropogenic processes have led to an increase in atmospheric N2O from 285 ppbv to about 320 ppbv, and an increase in the amount of reactive nitrogen [nitrate (NO3-), ammonium (NH4+), and other oxidized and reduced N species] circulating in the land-ocean system. Some estimates indicate that the amount of anthropogenic reactive nitrogen of 10 Tmol N/yr (140 Tg N/yr) is roughly equivalent to the preindustrial flux; that is, reactive nitrogen flow in the early 21st century is approximately double what it was in preindustrial time (Vitousek, Mooney, Lubchenco, & Melillo, 1997; Galloway et al., 2003; Mackenzie & Lerman, 2006).
The development of ammonium synthesis via the Haber-Bosch reaction in the early 20th century, which allowed for the production of inorganic nitrogen by the chemical industry, had a significant impact on the nitrogen cycle. The Haber-Bosch process became the dominant source of additional nitrogen input to the ecosphere, with additional input from fossil fuels and cultivation-induced biological nitrogen fixation. The characteristic signature of anthropogenic nitrogen production is that it tends to follow global population trends from 1860 to the early twenty-first century, with a major inflection downward in the early 1990s due to economic factors, and then a return to steadily increasing production. Although much of the global population has been fed by the fertilization and cultivation of crops, the negative impacts of excess nitrogen in the environment have led to issues of tropospheric air quality, destruction of the stratospheric ozone layer, a decrease in biodiversity, acidification of freshwater systems, and eutrophication and degradation of coastal ecosystems.
The geologic and modern biogeochemical cycles of P are distinguished by the fact that there is no significant transport of P as a gas phase in its global cycle. Phosphine (swamp gas, “will-o’-the-wisp”) is s gaseous source of P to the atmosphere but it is a minor flux in terms of other P fluxes related to the weathering of organic matter and inorganic phases in rocks. The latter are found in rocks and sediments as the minerals apatite and carbonate fluoroapatite. In addition, iron oxyhydroxides in soils sequester P in their structures. Due to chemical weathering, all these inorganic phases, along with organic matter, are sources of P to rivers and lakes and ultimately to the ocean. In the oceanic sediments, P is buried in the form of organic matter, iron oxyhydroxides, and carbonate fluorapatite. With deep burial, much of the P is converted to carbonate fluorapatite and on uplift, along with igneous apatite, because of chemical weathering, becomes a source of P for aquatic systems.
The global P cycle has been perturbed by human activities, with about 6.1 Tmol P/yr (190 Tg P/yr) added to the landscape and aquatic systems. The major source of this P is the mining of phosphorus-rich rocks, with most of the reserves found in Morocco, China, South Africa, and the United States. These countries hold 83% of the world’s easily exploitable phosphate rock, and account for two thirds of annual phosphorus production. It has been estimated that globally 90 years of the phosphorus resource remains (Vaccari, 2009). Ninety percent of the world’s phosphate rock is used as fertilizer in agriculture and food production. Minor amounts are consumed in animal feed and used as food additives. Land-use changes release phosphorus to soil water and continental rivers through processes of biomass burning and remobilization and erosion of P in soils and rocks. Since P, along with N and potassium (K), is an important nutrient for crops, it is also supplied to the land and continental waters in significant amounts in the form of P-bearing fertilizers. P is also released to aquatic systems through detergent use and sewage disposal. One outcome of these releases of anthropogenic P (and N) to the environment is the eutrophication of lacustrine and coastal marine waters, particularly in developing countries. It is likely that human-induced sources of P (and N) in the near future will continue to increase and areas of aquatic hypoxia and anoxia increase.
In the ocean, a complex set of processes governs the biogeochemical cycles of N and P. In general, N and P are taken up in the photic zone by phytoplankton. Zooplankton feed on the phytoplankton to accumulate N and P and also release the elements to the water column in their feces. On death of the plankton, much of the organic N and P is remineralized and released to the shallow-water column to be available once more for bioproductivity. However, some particulate organic matter sinks out of the photic zone into the deeper ocean carrying with it N and P. A substantial portion of particulate organic matter is decomposed in the deeper ocean, where the N and P are released to the water column as dissolved inorganic and organic forms of N and P. Upwelling waters can carry this dissolved N and P back to the surface waters, where the elements can be utilized once more in phytoplankton productivity. Some N and P in particulate organic matter sediments out of the ocean and is buried in marine sediments, along with river-borne detrital N and P in organic matter, and, in the case of P, as iron oxyhydroxide and apatite phases.
The relative role of N and P in globally limiting Earth’s bioproductivity is still a matter of debate. This is a highly relevant issue in the current context of the increasing eutrophication observed in lakes, estuaries, and coastal waters of the ocean. However, from a global perspective, the C:N:P ratio in biota and the relative abundance of N and P in the environment suggest that P is the limiting nutrient. In the ocean, P is mainly furnished by river discharge, aeolian inputs, and the remineralization of organic-rich sediments. On land, P is supplied to soil water through microbial degradation of organic matter and by the weathering of P-bearing minerals. This contrasts with the sources of N on land: although weathering does deliver N, mainly as nitrate to soil water for use in organic production, large amounts of atmospheric N can be taken up by leguminous plants through the biotic fixation of N2. In the ocean, atmospheric N also represents an important N source for living organisms that directly fix atmospheric N2, such as the blue-green algae (Cyanobacteria). Finally, the relatively low solubility of P-bearing minerals, such as apatite and Fe- and Al-phosphates, also suggests a limiting role of P in the growth of land biota and thus is a significant influence on the global carbon cycle.
Coupling of the Carbon, Nitrogen, and Phosphorus Global Biogeochemical Cycles
As knowledge about the global biogeochemical cycles of the elements has accumulated, there have been increasing attempts to develop more sophisticated conceptual and numerical models of the cycles, particularly in an attempt to understand the behavior of the cycles and predict the influence of human activities on the biogeochemical cycles and their ties to environmental problems. An example of one such attempt is that of the coupling of the cycles of the bioessential elements C, N, and P. Figure 11 is a conceptual multireservoir diagram of the coupled global biogeochemical cycles of carbon, nitrogen, and phosphorus illustrating the processes responsible for the movement of these elements between reservoirs of the Earth’s surface system or outer shell and their associated fluxes. The major domains of C, N, and P are the land, coastal ocean, open ocean, sediments, and, in the case of C and N, the atmosphere. The coupling of the C, N, and P cycles is founded on the occurrence of these elements as the major constituents of organic matter, where atomic abundance ratios of C, N, and P are characteristic of terrestrial and aquatic primary producers. The ratios of these elements in organic matter have been termed Redfield ratios. For example, the average ratio of C:N:P in marine phytoplantkon is 106:16:1 (Figure 11), and this ratio is maintained in calculations using a numerical model based on the conceptual diagram TOTEM (Terrestrial-Ocean-aTmosphere Ecosystem Model). TOTEM is a 16-box (reservoir) model (see Section II) with processes and fluxes representative of the steady-state flows of C, N, and P in late preindustrial time before strong human interference in the global cycles of these elements, as seen in the diagram of Figure 11.
TOTEM has been used to make projections of the behavior of the C cycle and its driver nutrients of N and P into the early decades of the 21st century. The initial steady state of TOTEM (Figure 11) has been perturbed using the important anthropogenic forcings of (1) atmospheric CO2 emissions from fossil-fuel burning, land -use activities, and cement manufacturing; (2) atmospheric gaseous NOx and SOx emissions mainly from fossil-fuel combustion; (3) the application of N and P-bearing fertilizers to croplands; (4) C, N, and P sewage inputs to aquatic systems; and (5) changes in global mean temperature. The model calculations of the rise in atmospheric CO2 concentrations over 300 years of the Anthropocene due to fossil-fuel and land-use change emissions are in very good agreement with the historical observational CO2 record from ice cores and more recently from Mauna Loa (see Ver, Mackenzie, & Lerman, 1999; Mackenzie, Ver, & Lerman, 2001, for details).
One question a global biogeochemical cycle model can answer is: Where do all the anthropogenic carbon, nitrogen, and phosphorus go after being emitted to the environment by human activities? Figures 12 and 13 show the partitioning of the human-produced perturbation fluxes on land as calculated from TOTEM for C, N, and P for the period 1850 to the year 2040 under a business as usual projection for anthropogenic sources of the elements from the early twenty-first century to 2040. The strength of the anthropogenic sources of the elements are plotted below the zero line (negative values) and resulting accumulation and enhanced export fluxes above the zero line (positive values). The diagrams balance only for the anthropogenically derived fluxes; in the background, there are large exchanges of materials that go on naturally as represented in Figure 11.
The sinks for anthropogenic CO2 emissions since 1850 to the early twenty-first century with projections to 2040 are shown in Figure 12 as positive values above the zero line. During the past 300 years, a cumulative amount of carbon from fossil-fuel combustion and changing land-use patterns of approximately 41.7 × 1015 mol (500 Gt C) were emitted to the atmosphere, This anthropogenic carbon came to reside in the sinks of the atmosphere, leading to most of the rise in atmospheric CO2 concentrations during this time period, ocean, terrestrial biotic uptake, and less importantly in the accumulation of organic matter in coastal marine sediments. Because of the release of about 16.7 × 1015 mol (200 Gt C) during this 300-year time period due to land-use activities and increased river discharges of organic and inorganic carbon and sewage discharge of organic carbon to the ocean, the land actually lost a cumulative net amount of carbon of approximately 5.8 × 1015 mol (70 Gt) (e.g., Lerman, Mackenzie, & Ver, 2004).
In concert with the CO2 emissions from the land for the period 1850 to the early twenty-first century, there were major anthropogenic inputs of the nutrients N and P to the landscape. For N the major inputs were from atmospheric deposition, leaching of fertilizer, and land-use remobilization of N from humus and the living terrestrial phytomass. These fluxes were balanced by N accumulations in the biomass and humus, export of dissolved and particulate N to the coastal oceans, and loss of N via denitrification and emission of gaseous N forms (Figure 13A). It also can be seen from Figure 13(A) that the anthropogenic mobilization of N in the Earth’s surface environment became dramatically more important following World War II as a result of the growing human population and its demand for resources to fuel the expanding global economy of the postwar years. The total N mobilized on the Earth’s surface by human activities was more than three times greater than that at the close of World War II, when much of the mobilized N was associated with land-use activities (about 90%). As the twentieth century progressed, proportionately more of the anthropogenic N emissions came from fertilizers and the combustion of fossil fuels. In the Earth surface system, the anthropogenic fluxes were redistributed into the environment by several processes. The application of N fertilizers to the landscape resulted in the leaching of some of the nitrogen into aquatic systems as dissolved NO3− and NH4+ and an increase in the export of dissolved inorganic nitrogen (DIN) and total organic nitrogen (TON) to the coastal ocean. The application of nitrogenous fertilizers to croplands also led to an increase in the export flux of N to the atmosphere, including the emissions of N2 and N2O gases by the process of denitrification. Finally, it is also of interest that fertilizer N remaining in the soil water, N from atmospheric deposition, and N remobilized by land-use activities via the degradation of living phytomass and humus probably led to the enhanced fertilization of the terrestrial phytomass, particularly some forested areas of temperate North America and Europe, and hence sequestration of atmospheric CO2 in phytomass.
The case for phosphorus shown in Figure 13(B) is similar to that for N with one very important exception: there is no major global flux of P as a gaseous species. Also, dust transport of P, mainly bond in iron oxyhydroxides, from the land to the ocean is a moderately important flux for getting P to the ocean but not for N. On a global scale, humans have significantly perturbed the P cycle by the mobilization of nutrient P through land-use activities and the application of phosphate fertilizer to croplands. For the past 150 years, changing land-use practices have increased the mobilization of P into the environment by a factor of about five and also increased loss of P from land due to erosion by a factor of 7. Additionally, the increased application of phosphate fertilizers, particularly since the close of World War II, provided a greatly enhanced source of anthropogenic P in the environment. The leaching of P from P-bearing fertilizers represented about 40% of the total P-bearing fertilizer applied annually in the early 21st century to croplands, golf courses, home lots, and so forth. Phosphorus leaching from these sources subsequently enters aquatic systems as dissolved inorganic and organic P (DIP, DOP) and as particulate P (PP). A portion of this P is exported to the coastal marine realm by river and groundwater discharges, while the rest is used in the fertilization of new phytomass. The P exported to aquatic systems can contribute to the problem of cultural eutrophication. Recently there has been concern voiced that the world is running out of easily exploitable P mined from P-bearing rocks and guano (Vaccari, 2009).
Our Earth’s basic metabolism is influenced by the circulation of numerous chemical elements and compounds through its atmosphere, soils and rocks, terrestrial and marine phytomasses, the flowing waters of rivers, lakes and oceans, and icy terrains in the form of global biogeochemical cycles. The circulation is driven ultimately by the sun and to, a lesser extent, by internal heat from the Earth. The global biogeochemical cycles of the elements and compounds operate on various time scales from that of millions of years to that of human generations. Processes and feedbacks within the cycles control, for example, the concentration of CO2 in the atmosphere; element concentrations in the ocean; natural nutrient availability in worldwide biological communities; and the discharge of water, suspended sediments, and dissolved substances to the ocean. Although the processes and fluxes associated with the global biogeochemical cycles are fundamentally natural, they are being strongly influenced and interfered with by the activities of human civilization, such as the combustion of coal, oil, and gas; deforestation; the application of fertilizers, pesticides, and herbicides to croplands; sulfur and nitrogen emissions from burning fossil fuels, and so forth. These activities of humanity have led to additional fluxes entering the biogeochemical cycles and a plethora of environmental problems.
The global biogeochemical cycles of the three most important bioessential elements of C, N, and P have been greatly affected by human activities, so much so that during the Anthropocene, significant environmental disruptions have occurred; for example, the radiative balance of the planet has been modified as a result of changing atmospheric composition due to the emissions of anthropogenic greenhouse gases to the atmosphere, and worldwide aquatic systems have become eutrophicated due to mining of phosphate ores and the production of fixed N for the manufacturing of fertilizers. This article has shown the fate of the sources and sinks of the anthropogenic C, N, and P added to the ecosphere. The addition of C as CO2 to the atmosphere is the principle cause of warming of the planet in the Anthropocene. In addition, excess organic carbon has been produced by human activities and transported into aquatic systems. The application of N- and P-bearing fertilizer to croplands and other parts of the landscape have led, through leaching, to excess dissolved N and P inputs into aquatic systems and increased biological productivity and production of organic matter. Coastal marine ecosystems worldwide have become hypoxic to anoxic because of the consumption of oxygen by microbially-mediated oxidation of the excess allochthonous organic carbon and the autochthonous organic carbon produced from the excess nutrient supply. To mitigate these environmental problems, it will be necessary to diminish greatly the magnitude of the fluxes added to the natural global biogeochemical cycles of the elements. To do less will mean increased global warming and degradation of our aquatic systems with consequent worldwide effects on biological community structure and biodiversity.
Andersson, A. J., Bates, N. R., & Mackenzie, F. T. (2007). Dissolution of carbonate sediments under rising pCO2 and ocean acidification: Observations from Devil’s Hole, Bermuda. Aquatic Geochemistry, 13(3), 237–264.Find this resource:
Andersson, A. J., & Mackenzie, F. T. (2004). Shallow-water oceans: a source or sink of atmospheric CO2? Frontiers in Ecology and the Environment, 2(7), 348–353.Find this resource:
Andersson, A. J., Mackenzie, F. T., & Bates, N. R. (2008). Life on the margin: implications of ocean acidification on Mg-calcite, high latitude and cold-water marine calcifiers. Marine Ecology Progress Series, 373, 265–273.Find this resource:
Arrhenius, S. (1896). On the influence of carbonic acid in the air upon the temperature of the ground. Philosophical Magazine, 41, 237–276.Find this resource:
Arvidson, R. S., Mackenzie, F. T., & Guidry, M. W. (2006). MAGic: A Phanerozoic model for the geochemical cycling of major rock-forming components. American Journal of Science, 306(3), 135–190.Find this resource:
Arvidson, R. S., Mackenzie, F. T., & Guidry, M. W. (2013). Geologic history of seawater: A MAGic approach to carbon chemistry and ocean ventilation. Chemical Geology, 362, 287–304.Find this resource:
Berner, R. A., & Krothavala. (2001). Geocarb III: A revised model of atmopsheric CO2 over Phanerozoic time. American Journal of Science, 301, 182–204.Find this resource:
Burchfield, J. D. (1975). Lord Kelvin and the age of the Earth. New York: Science History Publications.Find this resource:
Crawford, E. (1997). Arrhenius and the greenhouse gases. Ambio, 26(1), 6–11.Find this resource:
Doney, S. (2006, May). The dangers of ocean acidification. Scientific American, 294, 58–65.Find this resource:
Drupp, P. S., De Carlo, E. D., & Mackenzie, F. T. (2016). Porewater CO2-carbonic acid system chemistry in permeable carbonate sands. Marine Chemistry, 185, 48–64.Find this resource:
Drupp, P. S., De Carlo, E. H., Mackenzie, F. T., Sabine, C. L., Feely, R. A., & Shamberger, K. E. (2013). Comparison of CO2 dynamics and air-sea exchange in differing tropical reef environments. Aquatic Geochemistry, 19(5–6), 371–397.Find this resource:
Ekholm, N. (1901). On the variations of the climate of the geological and historical past and their causes. Quarterly Journal of the Royal Meteorological Society, 27, 61.Find this resource:
Feely, R. A., Sabine, C. L., Hernandez-Ayon, J. M., Ianson, D., & Hales, B. (2008). Evidence for upwelling of corrosive “acidified” water onto the continental shelf. Science, 320, 1490–1492.Find this resource:
Fleming, J. R. (2001). The carbon dioxide theory of climate change: emergence, eclipse, and reemergence, ca. 1850–1950.
Fleming, J. R. (2005). Historical perspectives on climate change. New York: Oxford University Press, p. 208.Find this resource:
Galloway, J. N., Aber, J. D., Erisman, J. W., Seitzinger, S. P., Howarth, R. W., Cowling, E. B., et al. (2003). The nitrogen cascade. Bioscience, 53, 341–356.Find this resource:
Garrels, R. M., & Mackenzie, F. T. (1971). Evolution of sedimentary rocks. New York: W. W. Norton.Find this resource:
Garrels, R. M., & Mackenzie, F. T. (1972). A quantitative model for the sedimentary rock cycle. Journal of Marine Chemistry, 1, 27–40.Find this resource:
Gattuso, J.‑P., Frankignoulle, M., Bourge, I., Romaine, S., & Buddemeier, R. W. (1998). Effect of calcium carbonate saturation of seawater on coral calcification. Global and Planetary Change, 18(1–2), 37–46.Find this resource:
Gattuso J.‑P., & Hansson, L. (2011). Ocean acidification. Oxford: Oxford University Press.Find this resource:
Gregor, C. B., Garrels, R. M., Mackenzie, F. T., & Maynard, J. B. (1988). Chemical Cycles in the Evolution of the Earth. New York: Wiley.Find this resource:
Intergovernmental Panel on Climate Change. (2007). Climate Change 2007: The Physical Science Basis. New York: Cambridge University Press.Find this resource:
Intergovernmental Panel on Climate Change. (2013). Climate Change 2013: The Physical Science Basis. New York: Cambridge University Press.Find this resource:
Jokiel, P. L., Rodgers, K. S., Kuffner, I. B., Andersson, A. J., Cox, E. F., & Mackenzie, F. T. (2008). Ocean acidification and calcifying reef organisms: A mesocosm investigation. Coral Reefs, 27, 473–483.Find this resource:
Joly, J. (1899). An estimate of the geological age of the Earth. Royal Dublin Society, 7, 23–66.Find this resource:
Kleypas, J. A., McManus, J. W., & Menez, L. A. B. (1999). Environmental limits to coral reef development: Where do we draw the line? American Zoologist, 39, 146–159.Find this resource:
Kuffner, I. B., Andersson, A. J., Jokiel, P. L., Rodgers, K. S., & Mackenzie, F. T. (2008). Decreased abundance of crustose coralline algae due to ocean acidification. Nature Geosciences, 1, 114–117.Find this resource:
Le Quéré, C., Moriarty, R., Andrew, R. M., & Zeng, N. (2015). Global carbon budget 2015. Earth System Science Data, 7(2), 349–396.Find this resource:
Lerman, A., Mackenzie, F. T., & Garrels, R. M. (1975). Modeling of geochemical cycles: Phosphorus as an example. In E. H. T. Whitten (Ed.), Quantitative Studies in the Geological Sciences: A memoir in honor of William C. Krumbein. Geological Society of America Memoir, 145, 205–218.Find this resource:
Lerman, A., Mackenzie, F. T., & Ver, L. M. (2004). Coupling of the perturbed C-N-P cycles in industrial time. Aquatic Geochemistry, 10, 3–32.Find this resource:
Lohbeck, K. T., Riebesell, U., & Reusch, T. B. H. (2012).Adaptive evolution of a key phytoplankton species to ocean acidification, Nature Geoscience, 5, 346–351.Find this resource:
Lohbeck, K. T., Riebesell, U., & Reusch, T. B. H. (2014). Gene expression changes in the coccolithophore Emiliana huxleyi after 500 generations of selection to ocean acidification. Proceedings of the Royal Society B, 281(1786), 1–7.Find this resource:
Lotka, A. (1925). Elements of physical biology. Baltimore: Williams and Wilkins.Find this resource:
Mackenzie, F. T. (1999). Global biogeochemical cycles and the physical climate system. Boulder, CO: University Center for Atmospheric Research, Global Change Instruction Program Module.Find this resource:
Mackenzie, F. T. (2011). Our changing planet: An introduction to earth system science and global environmental change. 4th ed. Upper Saddle River, NJ: Prentice Hall/Pearson.Find this resource:
Mackenzie, F. T., & Andersson, A. J. (2013). The marine carbon system and ocean acidification during Phanerozoic time. Geochemical Perspectives, 2(1), 1–227.Find this resource:
Mackenzie, F. T., & Lerman, A. (2006). Carbon in the geobiosphere: Earth’s outer shell. Dordrecht, The Netherlands: Springer.Find this resource:
Mackenzie, F. T., Ver, L. M., & Lerman, A. (2001). Recent past and future of the global carbon cycle. In L. C. Gerhard, W. E. Harrison, & B. M. Hanson & (Eds.), Geological perspectives of global climate change. Tulsa, OK: American Association Petroleum Geologists Special Publication, 47, 51–82.Find this resource:
Morse, J. W., Andersson, A. J., & Mackenzie, F. T. (2006). Initial responses of carbonate-rich shelf sediments to rising atmospheric pCO2 and ocean acidification: Role of high Mg-calcites. Geochimica et Cosmochimica Acta, 70, 5814–5830.Find this resource:
Odum, E. P., & Odum, H. T. (1959). Fundamentals of ecology. 2d ed. Philadelphia: W. B. Saunders.Find this resource:
Petit, J. R., Jouzel, J., Raynaud, D., Barkov, N. I., Barnola, J.-M., Basile, I., et al. (1999). Climate and atmospheric history of the past 420,000 years from the Vostok ice core, Antarctica. Nature, 399, 429–436Find this resource:
Pickett, M., & Andersson, A. J. (2015). Dissolution rates of biogenic carbonates in natural seawater at different pCO2 conditions: a laboratory study. Aquatic Geochemistry, 21(6), 459–485.Find this resource:
Rankama, K., & Sahama, T. G. (1950). Geochemistry. Chicago: University of Chicago Press.Find this resource:
Riebesell, U. (2004). Effects of CO2 enrichment on marine phytoplankton. Journal of Oceanography, 60, 719–729.Find this resource:
Regnier, P., Friedlingstein, P., Ciais, P., Mackenzie, F. T., et al. (2013). Anthropogenic perturbation of the carbon fluxes from land to ocean. Nature Geoscience, 6, 597–607.Find this resource:
Retallack, G. J. (2001) A 300-million-year record of atmospheric carbon dioxide from fossil plant cuticles. Nature, 411, 287–290.Find this resource:
Smith, S. V. (1972). Production of calcium carbonate on the mainland shelf of Southern California. Limnology and Oceanography, 17, 28–41.Find this resource:
Smith, S. V., & Mackenzie, F. T. (2015). The role of CaCO3 reactions in the contemporary oceanic CO2 system. Aquatic Geochemistry, 22(2), 153–175.Find this resource:
Tripati, A. K., Christopher, D. R, & Robert, A. E. (2009). Coupling of CO2 and ice sheet stability over major climate transitions of the last 20 million years. Science, 326, 1394–1397.Find this resource:
Vaccari, D. A. (2009, June). Phosphorus: A looming crisis. Scientific American, 300, 54–59.Find this resource:
Ver, L. M., Mackenzie, F. T., & Lerman, A. (1999). Biogeochemical responses of the carbon cycle to natural and human perturbations: past, present, and future. American Journal of Science, 299, 762–801.Find this resource:
Vernadsky, V. I. (1945). The biosphere and noosphere. Scientific American, 33, 1–12.Find this resource:
Vernadsky, V. I. (2007). Geochemistry and the biosphere. (F. B. Salisbury, Ed.). Santa Fe, NM: Synergetic Press.Find this resource:
Veizer, J. (1988). The evolving exogenic cycle. In C. B. Gregor, R. M. Garrels, F. T. Mackenzie, & J. B. Maynard (Eds.), Chemical Cycles in the Evolution of the Earth. New York: Wiley.Find this resource:
Vitousek, P. M., Mooney, H. A., Lubchenco, J., & Melillo, J. M. (1997). Human domination of Earth’s ecosystems. Science, 277, 494–499.Find this resource:
Walter, L., & Morse, J. W. (1985). The dissolution kinetics of shallow marine carbonates in seawater: A laboratory study. Geochimica et Cosmochimica Acta, 49(7), 1503–1513.Find this resource:
Woodwell, G., Mackenzie, F. T., Houghton, R. A., Apps, M., Gorham, E., & Davidson, E. (1998). Biotic feedbacks in the warming of the earth. Climatic Change, 40, 495–518.Find this resource:
Yapp, C. J., & Poths, H. (1992). Ancient atmospheric CO2 pressures inferred from natural goethites. Nature, 355, 342–344.Find this resource: