SCOPE 13 - The Global Carbon Cycle

ABSTRACT

The biogeochemical cycle of carbon constitutes the basic mechanism for the production of renewable resources such as food, fibre, and fuel and for the removal of organic waste material through mineralization. Man is increasingly affecting the carbon cycle by combustion of fossil fuels, by intensifying agriculture, and by destroying segments of the vegetation cover of the earth. In particular, the increasing CO₂ content of the atmosphere causes grave concern because of the climatic changes that may result from it. Thus, while having to rely to an increasing extent of renewable resources, man at the same time is now modifying in an as yet unpredictable manner the very process that is supplying these renewable resources.

In this introductory chapter, an atmospheric researcher (B. Bolin), a chemical oceanographer (E. T. Degens), an ecologist (P. Duvigneaud), and a geologist (S. Kempe) have pooled their talents: (i) to assess the airwaterliferock interactions as they determine the global biogeochemical cycling of carbon; and (ii) to define problem areas in the field of carbon where man's impact will be felt in the immediate future. The chapter also serves as a springboard for more detailed discussions on special topics presented in the follow-up chapters of the book.

1.1 INTRODUCTION

Carbon is the key element of life. The carbon cycle is, therefore, of fundamental importance when trying to understand the biosphere and its basic mechanisms and, conversely, the latter play an important part in determining the characteristics of the carbon cycle. The availability of carbon as carbon dioxide in the air, as carbonates in the earth's crust, as carbonate ions in the sea, and in the many organic compounds in land biota, the soil, and in the sea is basically dependent on the fact that gases containing carbon (methane and carbon dioxide primarily) escaped from the interior of the earth during geologic ages. The biosphere, as it exists today, has evolved in a complex interplay between carbon and many other elements; primarily hydrogen, oxygen, the basic nutrient elements, nitrogen, phosphorus, sulphur, and some metals in minor quantities that are fundamental to the development of life. It is for this reason that the carbon cycle cannot be dealt with independently of the cycles of other elements involved in the biogeochemical system.

A treatment of the carbon cycle implies consideration of physical, chemical, biological and geological processes that proceed on very different time scales, that is, from millions of years for the slow movement of the earth's crust, to weeks and days for the rapidly changing scene of the water surface of the sea. The slow processes are determined by the time required to accomplish significant changes in the earth's crust, for instance. It is of note, however, that in geology the so-called 'slow processes' frequently represent extended periods of quiescence which are briefly interrupted by active pulses of relatively short duration. In contrast, rapid transfers basically depend on the characteristic motions of air and water. We know, for example, that the most slowly varying component of the hydrologicalcycle is deep ground-water, the amount of which is small, and the cryosphere, i.e. ice and snow, the time scale of which primarily depends on the evolution of the Antarctic and Greenland ice sheets, i.e. times of at least 10 000 to 100 000 years. We shall also assume here these features to be constant. The characteristic turnover time of the oceans, about 1000 years, becomes the longest time scale to be dealt with explicitly in this introductory chapter.

Let us first focus our attention on the comparatively rapid processes in the carbon cycle. We need a thorough understanding of the quasi-balance that they maintain, in order to appreciate how the biosphere prevails and shows only slow global changes. The stability of the biosphere is to a considerable extent dependent on how the rather rapid processes interact. These problems are of course fundamental in themselves, but are of particular relevance today, as man has become a prime biogeochemical factor, perturbing the natural balance. Man's intervention consists primarily in the burning of fossil fuels and the changing of land use. The observed changes in the atmospheric concentration of carbon dioxide show directly that these interventions are significant on a global scale and simple model computations indicate that very marked changes, with possibly far-reaching consequences for the climate of the earth, may take place during the next century. A far better understanding of the key processes that control the carbon cycle is, therefore, of great importance if we are to be able to foresee the consequences of man's continued interference with the carbon cycle on a global scale.

In later chapters of this book, detailed accounts are given of the carbon cycle subsystems and their characteristics, including processes over geological time scales. We shall here briefly summarize those features that will be needed to build a dynamic model of the carbon cycle. In order to deal with the problem quantitatively, rather drastic simplifications are needed, which must be justified as far as possible. In doubtful cases, the various possibilities of dealing with the problem should be explicitly modelled to resolve ambiguities.

The modelling approach was begun about twenty years ago (Craig, 1957; Revelle and Suess, 1957; Bolin and Eriksson, 1959). Many important developments have been published since then (Broecker et al., 1971; Machta, 1972; Keeling, 1973a; Oeschger et al., 1975; Bolin, 1975; Keeling and Bacastow, 1977). Keeling's article (1973a) is particularly useful in the present context as we wish to synthesize our knowledge into one general model. The following presentation is, to a considerable extent, based on his results. A more detailed treatment of the ocean circulation and the land biota is presented by Björkström (see Chapter 15, this volume). A recent article by Revelle and Munk (1977) is also of considerable interest. Additional light will be cast on the question of the relative importance of oceans and the land biota as the most important sink for atmospheric carbon dioxide. This is a crucial question if we wish to analyse the observed changes in the atmosphere in relation to man's emissions of carbon dioxide by burning fossil fuels, deforestation and increasing the land areas used for agriculture.

1.2 MAJOR RESERVOIRS

A schematic diagram, showing the main global reservoirs and fluxes, is depicted in Figure 1.1. Together with Table 1.1 it should serve as a reference for the presentation in this section.

1.2.1 The Atmosphere 

A. Total Concentration

The present (1977) concentration of carbon dioxide in the atmosphere is about 329 ppm (per volume). Since accurate measurements began in 1957 (Keeling and Bacastow, 1977; Bischof, 1977) an increase of about 17 ppm has been recorded. This result is based on measurements at Mauna Loa in Hawaii, at the South Pole, and from aircraft flights at middle- and high latitudes in the northern hemisphere (see Figure 1.2 and Chapter 3, this volume).

It has long been assumed that the preindustrial CO₂ level, i.e. before 1850, was between 290 and 295 ppm, which value has been determined by:

1. extrapolation back in time, based on the fact that the observed increase of carbon dioxide in the atmosphere 195775 has been 54% of the emissions due to fossil fuel combustion (Keeling and Bacastow, 1977), that this ratio was the same from 1850 to 1957 and that the use of fossil fuels was that estimated by Keeling (1973b); and

2. an assessment of the reliability of measurements during the latter part of last century (Bray, 1959) and the determination of the most likely value.

We note, however, that man has also been emitting carbon dioxide by deforestation and expansion of arable land (Bolin, 1977a; Revelle and Munk, 1977; Bohn, 1978; Woodwell and Houghton, 1977; Woodwell et al., 1978; see also Chapter 4, this volume) and that the measurements in the middle of last century were quite unreliable. The atmospheric carbon dioxide concentration may, therefore, have been lower and the degree of uncertainty is most likely at least 10 ppm.

Figure 1.1 Principal reservoirs and fluxes in the carbon cycle (WOW = warm ocean water; COW = cold ocean water; POM = particulate organic matter; DOM = dissolved organic matter)

Table 1.1 Major carbon reservoirs and fluxes

Reservoirs			1015g C
Atmosphere: Before 1850
	Common assumption:	290 ppm	610
	Stuiver (1978)	265 ppm	560
	Chapter 15, this volume
	1978	329 ppm	692
Oceans: Total amounts, inorganic			35 000
	Above thermoclinie, low and middle latitudes		600
	Dissolved organic matter, Chapter 11, this volume		1 000
	Particulate organic matter, biomass, Chapter 10, this volume		3
Land biota: Whittaker and Likens (1975)			827
	Bazilevich et al. (1970)		976
	Duvigneaud, this chapter		592
Soil, humus: Keeling (1973a)			1 050
	Bohn, 1978		3 000
	Duvigneaud, this chapter		2 840
Sediments: Total			>10 000 000
	Available for dissolution in oceans
	Broecker and Takahashi (1977)		4 900
Fossil fuels			>5 000
Fluxes			1015g C/year
Atmosphere: oceans, gross exchange			100
Atmosphere: land biota, photosynthesis
	Whittaker and Likens (1975)		53
	Bazilevich et al. (1970)		78
	Duvigneaud, this Chapter		63
Ocean photosynthesis: Chapter 10, this volume
	maximum		126
	minimum		15
	adopted average		45
Lands to oceans: Chapter 12, this volume
	dissolved inorganic matter		0	.4
	dissolved organic matter		0	.1
	particulate organic matter		0	.06
Deposition in oceans: Peng et al. (1977)			1	10
Fossil fuel combustion, 1978			5

The rate of atmospheric CO₂ increase has not been constant during these twenty years of observation, as seen from Figure 1.2 (Keeling and Bacastow 1977; Bischof, 1977). The data from the South Pole and Mauna Loa show oscillations, probably associated with the climatic variations known as the Southern Oscillation, which is a quasi-periodic oscillation of the major wind systems over the Central and South Pacific.

Figure 1.2 Seasonally adjusted concentrations of atmospheric CO₂ at the Mauna Loa, South Pole (Keeling and Bacastow, 1977), and as measured North from commercial aircraft over the Polar region (Bischof, 1977). The North Polar curve is a second-degree polynomial deduced from average annual values; the values have not been corrected by intercalibration with the former data series

The atmospheric carbon dioxide concentration also shows marked seasonal variations (Bolin and Keeling, 1963; Bolin and Bischof, 1970; Keeling and Bacastow, 1977; Bischof, 1977). They are due primarily to changes of assimilation by land biota between winter and summer in the northern hemisphere. The variation with latitude and height of this seasonal variation may be interpreted in terms of large-scale atmospheric mixing and is consistent with a characteristic vertical exchange time in the troposphere of about one month, and a horizontal exchange time between the northern and southern hemispheres of about one year. It is well known that the mixing of the stratosphere is considerably slower. This is also obvious from the fact that the seasonal variations rapidly decrease as a function of height above the tropopause (Bischof, 1977). As we shall essentially be concerned with the evolution of the carbon cycle over decades, centuries and millenia, the atmosphere may, in most cases, be treated as well-mixed and we shall assign average annual values, which may change slowly from one year to the next.

B. Isotopic Composition

There are three carbon isotopes in nature:¹²C,¹³C, and ¹⁴ C, with the approximate relative amounts 1, 10^-2, and 10^-12. The former two are stable, while ¹⁴C is produced in the atmosphere by cosmic radiation reacting with nitrogen, and decays with a half-life of 5570 years. Since 1952, ¹⁴C has also been emitted into the atmosphere by nuclear bomb tests and from nuclear power plants.

Figure 1.3 ¹³ C record of trees from various regions. (a) Tasmanian King billy pine, cellulose (after Pearman et al., 1976). (b) English oak and larch, whole wood (after Farmer and Baxter, 1974). (c) Russian spruce, whole wood (?) (after Galimov, quoted in Rebello and Wagener, 1976). (d) Oak and pine, cellulose (triangles, Freyer and Wiesberg, 1974), (crosses, Rebello and Wagener, 1976) and Douglas Fir, cellulose (circles, Stuiver, 1978)

Figure 1.4 Total amount of ¹⁴C in the atmosphere and troposphere in excess of pre-bomb conditions (after Machta, 1972)

The ¹³C content of the atmosphere is 7‰ less than the PDB ¹³ C-standard, while the corresponding ¹³C value for modern wood and fossil fuel carbon is 25‰.

Analyses of ¹³C in tree rings show, however, a decrease of 11.5‰ over the past 200 years; see Figure 1.3 (Freyer and Wiesberg, 1974; Farmer and Baxter, 1974; Rebello and Wagener, 1976; Stuiver, 1978). Since the fractionation between ¹² C and ¹³ C during the assimilation process has remained the same, this change may be interpreted as a change of ¹³C in the air. The fractionation factor is slightly dependent on temperature, but this effect can be corrected for.

Similarly, a decrease (the 'Suess-effect') of about 2% was recorded in the ¹⁴C content of wood from the middle of the last century until 1954 (Keeling, 1973a). This is due to the dilution that is caused by emitting CO₂ by fossil fuel combustion, as it contains no ¹⁴C. Since 1954, large amounts of ¹⁴C have been added to the atmosphere because of nuclear bomb testing (Machta, 1972) (Figure 1.4), causing a marked increase in ¹⁴C concentration until 1962, at which time atmospheric testing almost ceased. Since that time, the ¹⁴C concentration has been gradually decreasing.

1.2.2 The Oceans

A. The Ocean Circulation

On the whole, the oceans are stably stratified and almost all the energy that maintains motions in the sea is supplied at the surface. Only when exceptionally dense water is formed at the sea surface by cooling or increased salinity when freezing occurs can it sink to appreciable depths. This occurs principally around the Antarctic and in a limited region in the IcelandGreenland area. These areas are the only sources for the very large volume of deep water in the world oceans. At the circumpolar convergence zone in the southern seas, and also on the polar side of the major ocean currents in the North Pacific and North Atlantic, the water sinks to intermediate depths (1001000 m) and moves towards the equator. Similarly, the rather warm, but saline Mediterranean water penetrates the Atlantic Ocean to intermediate levels. The sinking of cold water is balanced by very slow upward motions, which, at a first approximation, are evenly distributed over the entire ocean. Although the stratification is stable, some vertical mixing also occurs. The vertical distributions of ¹⁴C and oxygen have been used to assess the slow upward motion (w) and the vertical diffusivity (K), yielding values of w = 1 cm/day and K = 1 cm² /sec (Wyrtki, 1962).

B. Dissolved Carbon Compounds in the Sea

The concentration of inorganic carbon in sea water varies from between 1.95 mg-atoms/l in tropical surface water, to about 2.35 mg-atoms/l in the deep water. The total amount of inorganic carbon in the sea is about 35 000 x 10¹⁵ g. The ionic composition of sea-water is primarily determined by the carbonate and borate systems (see Bolin and Eriksson, 1959; Keeling, 1973a; and Chapter 9, this volume):

CO₂ +H₂O H⁺ + HCO₃

HCO₃ H⁺ + CO₃²

H₃ BO₃ H⁺ + H₂ BO₃

The amount of dissolved carbon dioxide in sea-water, which is the only species of the carbonate system in direct exchange with atmospheric CO₂, is only about one per mille of the total amount of dissolved inorganic carbon. Because of this fact and the necessary hydration and diffusion of the CO₂ molecules through a thin boundary layer at the sea surface, the rate of transfer and exchange between the atmosphere and the sea is limited (Bolin, 1960). Therefore, the turnover time between the atmosphere and the surface layers of the sea is 58 years. Quinn and Otto (1971) have pointed out that a more detailed treatment of the diffusive layer at the sea surface is necessary to describe the exchange processes properly, but the results of Bolin (1960) are on the whole still valid as quoted. This turnover time for atmospheric CO₂, relative to the ocean, agrees well with the observed decline in the ¹⁴C content of atmospheric CO₂ since 1962 (Figure 1.4). For a more detailed discussion, the reader should refer to Machta (1972).

The dissociation of sea-water, ion composition, and the amount of dissolved CO₂ gas, are dependent on the total amount of inorganic carbon in the sea-water, and the CO₂ partial pressure increases much more rapidly than the total amount of inorganic carbon. This ratio, called the buffer factor, is approximately 9 for a carbonateborate solution such as sea-water and at the present atmospheric partial pressure of 330 ppm (see Chapter 9, this volume). It approximately doubles if the partial pressure is doubled (Keeling, 1973a). This fact makes the oceans much less of a potential sink for the CO₂ emissions into the atmosphere than one would expect from the large size of the oceanic carbon pool.

In addition to the inorganic carbon, the oceans contain about 1000 x 10¹⁵ g of organic carbon (Chapter 11, this volume). This organic material is the decay product of life in the sea and also, to some extent, of the terrestrial biota carried to the ocean by rivers. An understanding of the turnover processes requires knowledge of the chemical composition of the dissolved organic matter and a more detailed study than has been done so far.

C. ¹⁴C in Sea-water

The radioactive carbon isotope ¹⁴C, contained in some atmospheric CO₂ molecules, is transferred to the sea. Due to the different molecular weight of ¹²CO₂ and ¹⁴CO₂, the transfer rates are not the same for the two molecules. This fractionation results in a ¹⁴C/¹²C ratio in surface water which is 0.95 of that of atmospheric CO₂. The ¹⁴C-activity is, therefore, often expressed as a deviation (in per mille) from 0.95 of a standard activity (Östlund et al., 1974), denoted by ¹⁴C. Due to the comparatively slow turnover of the sea and radioactive decay the ¹⁴C/¹²C C ratio of deep water is considerably lower, Figure 1.5 shows some characteristic vertical ¹⁴C profiles from the Atlantic and Pacific Oceans. The lowest value, 230 per mille, corresponds to an age of about 1800 years. Surface waters have been significantly influenced by the transfer of bomb-produced ¹⁴C to the sea since 1954. Positive ¹⁴ C-values have been measured down to layers below 400 m, particularly in the North Atlantic, which indicates that the influence of bomb-produced ¹⁴C may extend to these deeper layers. Small changes have also been observed in the Pacific Ocean (Broecker et al., 1978).

Figure 1.5 Vertical profiles of radiocarbon in sea-water in various regions in the Atlantic Ocean and in the Pacific Ocean. The data points are averages of a number of measurements as indicated. (Data from Broecker et al., 1960; Bien et al., 1960; Fairhall et al., 1972; and Östlund et al., 1974)

Östlund et al. (1974) also show conclusively that some ¹⁴C-rich water has penetrated to the bottom, in the area of deep-water formation in the North Atlantic.

The youngest deep water is found in the eastern North Atlantic, with the age of the water increasing southwards. Generally, Atlantic deep water is considerably younger than Pacific deep water where the age increases northwards (cf. Bien et al., 1960). These ages indicate the direction of flow of the deep water from the main formation areas in the North Atlantic and around the Antarctic continent.

A few ¹⁴C determinations of dissolved organic compounds are available (Williams et al., 1969). They  indicate an average ¹⁴C age of about 3000 years. Dissolved organic matter in surface water may be considerably younger, but little data is available.

D. Primary Productivity in the Ocean

Phytoplankton growth is influenced by a number of factors such as light intensity, availability of mineral nutrients, turbulent mixing, secondary production, and adaptation processes (see Chapter 10, this volume). In certain areas of the sea, the phytoplankton exhibits strong seasonal variability, the predominant feature being the outburst in spring. The global distribution of primary production is determined by the availability of nutrients (nitrogen and phosphorus) in the uppermost tens of metres, where there is sufficient light for effective photosynthesis. In areas of intense upwelling, the rate of photosynthesis reaches values of 5001000 g C/m² year, while in the 'desert' areas of slow sinking motion, where only small amounts of nutrients are available, the rate may be as little as 10% of this value or even less (see Figure 10.3, this volume).

For the following discussion, we shall adopt the values 45 x 10¹⁵ g C/year for the total primary production in the oceans and 3 x 10¹⁵ g C for the biomass. This yields a ratio of 15, corresponding to a turnover time of 24 days for living organic matter in the sea. For the development of a global carbon cycle model, we need to describe in some detail photosynthesis in the oceans and its dependence on environmental parameters. In doing so we should, however, keep the number of variables small. A first attempt of this kind was made by Riley et al. (1949) and model work has progressed ever since. Numerous field experiments have been carried out to obtain as complete descriptions as possible of the relevant processes, for example with regard to season (Steele, 1956, 1974), or diverse marine habitats (Riley, 1965). Reasonable agreement between theory and observation has been obtained to within one order of magnitude. Nevertheless, our understanding of the interrelationships affecting primary production remains limited. Progress is slow because of the intricate nature of life cycles and their reciprocal dependence on a multitude of environmental factors. As an illustration, some recent results from an experiment in the North Sea will be summarized, i.e. the Fladen Ground Experiment, FLEX 76. In other parts of the oceans with different climate and hydrography, the ecological interplay will, however, be quite different.

Fladen Ground Experiment, FLEX 76. To permit a detailed comparison between theory and a field experiment, a large number of environmental variables must be measured simultaneously with high spatial (in particular vertical) and temporal resolution. As a contribution in this regard, the Fladen Ground Experiment was carried out in 1976 (FLEX 76). This area was chosen because it is almost in the centre of the North Sea where horizontal gradients of physical conditions are relatively small and, therefore, the vertical exchange and transport processes are of primary importance for the balance. The depth in the Fladen Ground area is sufficient (150 m) to permit the development of a dynamically and biologically decoupled upper layer during the time of the experiment, but is also sufficiently shallow to permit the downward particulate flux, measured near the bottom, to be related to the process in the upper layer. An additional advantage of this site is that its general hydrography and productivity has previously been studied by Steele (1956, 1957a, b, c, d). Measurements of the physical, chemical, and biological processes were obtained in the entire water column. The spring plankton bloom was chosen for study. Full consideration was given to both horizontal and vertical variability.

The experiment took place within a 100-km-side square and covered a three-month period from mid-March to mid-June 1976. The initial phase concentrated on the analysis of factors which trigger the plankton bloom, and on observations of the rapid cell division which characterizes the onset of the bloom. The main phase of the experiment consisted of an intensive data acquisition programme lasting approximately one month. During the end phase, secondary programmes were carried out as a supplement to the main programme.

On account of the sensitivity of plankton growth rates, during the initial stages of an outburst, to small variations in the density structure and vertical exchange rates (Hasselmann, personal communication), a rather weak horizontal variability in these physical parameters can yield a pronounced patchiness of plankton and nutrient distributions. This makes it difficult to detect interactions between the horizontal variability of the physical and biologicalchemical components of the system by direct spectral correlation techniques.

To overcome some of these difficulties, a permanent station was occupied for the entire length of the experiment, while, at the same time, a drifting station followed a discrete `plankton patch' across the FLEX box.

Figure 1.6 Depth and time profiles of water temperature, phosphate, chlorophyll, Calanus finmarchicus, and colony-forming bacteria (CFU) at a central station of the FLEX Box positioned in the centre of the North Sea (58^°55´ N, 0^°32´ E); note that for graphical reasons the contour pattern for PO₄ concentration is light at high, and dark at low concentrations (after 'Meeresforschung in Hamburg'. Illustrated documentation of the `Sonderforschungsbereich 94' at the University of Hamburg, F.R.G., 1977)

Some critical data from the central station (58^°55´ N, 0^°32´ E) are depicted in Figure 1.6 to show the intricate relationships between water temperature, mineral nutrients, phytoplankton, zooplankton, and bacteria. Samples were taken at intervals of two to six hours; gaps in the profiles occur because of difficult ship logistics.

The principal factor controlling the onset of a plankton bloom appears to be related to the establishment of a thermocline. Such a phase boundary allows planktonic cells to stay in the euphotic zone for an extended period of time. As a consequence, the uptake of mineral nutrients by phytoplankton is accelerated, and, accordingly, the levels of dissolved minerals in Figure 1.6 exemplified by the phosphate curve are lowered in the upper layer. Following the first phytoplankton maximum (29 April to 2 May), grazers (a principal species is Calanus finmarchicus) and bacteria living on organic substances appear in larger numbers. The FLEX experiment clearly demonstrates that phase boundaries in mid-water have a pronounced effect on the pattern of dissolved mineral species. Mechanism and rate of molecular exchange across a well-developed pycnocline have been studied thoroughly (Craig, 1969; Spencer and Brewer, 1971; Brewer and Spencer, 1974). Vertical advection and diffusion will move deep water through the density boundary to the surface layer. For instance, in the main pycnocline in the Black Sea, at a water depth of about 200 m (Degens and Stoffers, 1976), the vertical eddy-diffusion coefficient is about 0.014 cm² /s (Brewer and Spencer, 1974). From this value, one can readily calculate the net upward flux of dissolved species such as hydrogen sulfide. Conversely, there is a downward diffusion gradient and transfer of, e.g., oxygen. Models using a constant eddy-diffusion coefficient for the vertical transfer of passive properties in a continuously layered medium (Riley et al., 1949) have only limited applicability because as we can see in the FLEX experiment density boundaries tend to rise and fall and many other physical perturbations also exist (Kraus and Turner, 1967; Denman, 1973).

E. Biomineralization

Photosynthesis transfers carbonate ions in the water to organic compounds and is thus an effective link in the carbon cycle. The bacterial decomposition of organic compounds returns carbon to the sea-water in dissolved form either as inorganic carbonate ions or organic dissolved compounds. Carbonate-forming organisms, in addition, remove carbonate from sea-water by extracting CaCO₃. This process of biomineralization is far more important than the inorganic precipitation of aragonite, calcite, and dolomite, which is globally insignificant.

Only recently has light been shed on the rather intricate biochemical process of biomineralization (Degens, 1976). In principle, carbonate deposition is an outgrowth of excretionary processes and related to the regulation of Ca²⁺ ions in the cytoplasm. The two main enzymes involved are an ATPase as Ca²⁺ transport enzyme and carbonic anhydrase; both have zinc ions at their active centre. The actual role of carbonic anhydrase in calcification does not depend on its ability to hydrate CO₂, but to remove carbonic acid from the site of calcification: 

Ca²⁺ + 2 HCO₃ Ca(HCO₃ )₂ CaCO₃ + H₂ CO₃

Figure 1.7 Electronmicrographs (a) taken in central region of pelecypod nacre showing steps formed by overlap of mineral laminae on growth surface; (b) vertical crosssection of pelecypod nacre exhibiting brick-wall pattern; (c) terminating screw dislocation (after Wise and de Villiers, 1971)

Thus, carbonate deposition involves interaction of Ca²⁺ and HCO₃, resulting in the formation of an unstable intermediary product Ca(HCO₃)₂. As long as calcium is not the limiting factor, the rate of formation of calcium carbonate will depend on the rate by which carbonic acid is removed from the calcification site. In the presence of carbonic anhydrase, calcification is significantly increased due to the formation of a complex between the high-pH form of the enzyme and a neutral H₂ CO₃ molecule, which is the substrate for carbonic anhydrase.

Respiratory CO₂ is often mentioned as the sole or principal source for the carbonate ion in biological calcite and aragonite. Stable isotope data, however, argue strongly against this supposition. Only in a few species can a significant contribution from respiratory CO₂ be demonstrated. Most invertebrates, and even the otoliths in fish, form their carbonate in, or almost in, isotopic equilibrium with the dissolved carbonate species present in their environment. In fact, the distribution of oxygen isotopes of the carbonates in many aquatic biological species reflects the water temperature in the habitat where the organisms live. Therefore, we must assume that bicarbonate enters the inner system of the cell via permeable membranes, where it forms a pool from which it is rapidly extracted, leaving little possibility for ¹³C/¹²C exchange with the isotopically lighter respiratory CO₂. Alternatively, carbonate deposition may proceed through direct contact with the outside environment.

The mediators in carbonate nucleation are organic templates with a strong affinity to calcium ions; they are composed of specific polysaccharides or proteins.

Figure 1.8 (a) A coccolithophorid alga, Emiliania huxleyi, and (b) deformed coccolith from the same species. The deformed specimen was collected from the North Central Pacific. All specimens of E. huxleyi collected from the area were malformed similarly to the one shown above (after Honjo, 1974)

A crucial point in the present discussion is the observation that biogenic carbonate extraction in the sea is chemically a non-equilibrium process, as far as the carbonate system is concerned. It is governed by a species-specific regulation mechanism involving a template mechanism which is based on solid-state principles. Inasmuch as carbonic anhydrase is the chief enzyme regulating carbonate formation, any inhibitors of carbonic anhydrase in the environment will affect rates of calcification. These inhibitors can be related to physical parameters such as abnormal temperature, salinity, certain pesticides, PCBs or nutrient deficiency. In Figure 1.8 a coccolithophorid alga in a normal and malformed stage is shown. Such severe malformation occurs when abnormal environmental conditions prevail, especially in the early stages of cell growth.

In summary, biomineralization is the principal mechanism of CaCO₃ extraction from the ocean. In view of the sensitivity of carbonic anhydrase activity to a number of environmental perturbations, care should be taken to reduce or eliminate any activity of man which inhibits carbonic anhydrase activity in carbonate depositing organisms.

F. Calcium Carbonate Dissolution and Ocean Sediments

The upper layers of the ocean are supersaturated relative to CaCO₃, while CaCO₃ starts to dissolve at greater depths due to the increasing hydrostatic pressure. The transition level is called the lysocline. At still greater depths, all calcareous remains exposed to sea-water completely dissolve at, and below, the carbonate compensation depth. This depth occurs around 3700 m. Inasmuch as we are dealing with a kinetic boundary, the actual depth varies as a function of type and quantity of infalling detritus or intensity of upwelling. For instance, in the equatorial zone of the Pacific, the CaCO₃ compensation depth may actually be positioned as much as 5000 m below the surface. On the whole, however, these transition levels lie considerably deeper in the Atlantic than in the Pacific Ocean. It is of interest to note that in areas where waters become stratified for an extended period of time, as is the case in the Black Sea or some deep East African Rift lakes, carbonate compensation depth is close to the thermo-halocline, i.e. at a water depth of 50 to 200 m. Carbonate particles that fall below this phase boundary dissolve, whereas those that settle in shallow areas above the pycnocline remain intact. Scanning electronmicrographs of chemically precipitated calcites from the Black Sea sedimented above and below the O₂ H₂S interface show this process (Figure 1.9).

Figure 1.9 Scanning electron micrographs of chemically precipitated calcites (a) sedimented above the Black Sea pycnocline and (b) below the pycnocline. Size of individual crystals 510 µm (after Degens and Stoffers, 1976. Reproduced by permission of Macmillan Journals Ltd. and by permission of the authors)

Figure 1.10 Schematic representation of the faecal pellet mechanism for a rapid removal of CaCO₃ from the surface to the bottom ocean. Sinking rates of coccoliths; in a pellet 160 m day; a discrete coccolith 0.15 m day (after Honjo, personal communication)

In view of the above relationship it has been difficult to explain why certain calcareous remains have escaped dissolution when settling below the critical CaCO₃ compensation depth. Organic coatings around individual mineral grains and organisms have been suggested as a critical inhibitory factor in CaCO₃ dissolution (e.g. Suess, 1973). Recently, Honjo (1977) has drawn attention to a transport mechanism, by means of faecal pellets. In principle, predators 'package' their excretion product which is full of calcareous remains in the form of pellets. These pellets may have a carbohydrate skin, a pellicle, or can be without a protective organic cover. Coccoliths may thus be rapidly transported from the euphotic zone, through the lysocline, to the deep ocean floor without dissolving. The sinking rate of a pellet is approximately 160 m per day compared to about 0.15 m per day for a discrete coccolith (Figure 1.10).

Obviously the characteristics of the bottom sediments are greatly influenced by these facts and we shall divide the ocean sediments accordingly as follows:

1. Those lying beneath the calcite compensation depth, where there are practically no calcium carbonate sediments. The only remains that are found are partly dissolved faecal pellets.

2. Those found between the calcite compensation depth and the lysocline, within which range there are some calcium carbonate sediments.

3. Those lying above the lysocline but below the continental shelves. These sediments are very rich in calcium carbonate, particularly at greater depth. 

4. Those on the continental shelves, where the deposits are a mixture of calcium carbonate and detritus from the continents.

Broecker and Takahashi (1977) have tried to estimate the amount of calcium carbonate with which ocean water could exchange calcium and carbonate ions. They consider that only about 10 cm of sediments would be available, which is equal to the mean burrowing depth of benthic organisms. Sediments below the bioturbated zone are protected from further dissolution. The amounts of carbonate available are 2400 x 10¹⁵ g C in the Atlantic and 1250 x 10¹⁵ g C in both Indian and Pacific Oceans. The total amounts of calcium carbonate sediments are, however, several orders of magnitude larger and could be in exchange with the sea-water during geological time scales.

For a dynamic treatment of the carbon cycle, we need to understand the processes that determine the rate of dissolution of the calcium carbonate from the sediment. The flux of calcium and carbonate ions from the carbonate grains in the sediment into the bulk ocean water is dependent on four distinct processes:

1. The dissolution at the interphase between the grains and the pore water;

2. the molecular diffusion in the pore water towards the sediment surface;

3. the molecular diffusion in a thin layer of water above the sediment surface; and

4. the turbulent transfer into the interior of the ocean which primarily occurs quasi-horizontally along density surfaces.

It is difficult to distinguish between 1 and 2 above, since this can only be done with certainty by detailed measurements in the pore water. Different models for this transfer in the top layers of the sediment have been developed, emphasizing one or the other process, the resaturation model or the stagnant-film model (Berner and Morse, 1974; Takahashi and Broecker, 1977). In the resaturation model, the rate-limiting process is that of CaCO₃ dissolution into the pore water. In the stagnant-film model, the molecular diffusion out of the sediment and the stagnant water film at the sediment surface are the rate-limiting processes. Takahashi and Broecker (1977) show that the observed decrease of sediment CaCO₃, with increasing depth below the lysocline, favours the former hypothesis with a characteristic time of 510 min for the rate-limiting process.

The formation of deep-sea carbonate sediments is essentially due to the fact that calcareous planktonic remains, which form in the photic zone, settle to the sea floor. The deposition rate can be determined from the ¹⁴ C-profile in the sediment at levels above the lysocline where no appreciable dissolution takes place (Peng et al., 1977). The sedimentation rates are in the range of 1 to 10 cm per 1000 years. In terms of carbon deposition, these values correspond to 0.55 x 10¹⁵ g C/year if integrated over the whole ocean surface, and a most probable value of 12 x 10¹⁵ g C/year.

1.2.3 Ecosystems on Land 

A. Basic Characteristics

It may seem simple to divide the biosphere into two major compartments, i.e. the oceans and the continents. The distinction is, however, not always clear. There is an extensive intertidal coastal zone of salt marshes, mangrove vegetation, sand beaches, and rocky shores. These ecosystems are among the most productive of the world. Similarly, fresh waters are intimately associated with land vegetation and soil processes.

Since the basic characteristics of the land vegetation is closely connected with the associated soil systems, it is necessary to deal with them simultaneously. This is particularly true when concerned with the dynamics of ecosystems during longer time periods (several decades or longer), for example, changes caused by climatic variations.

The natural ecosystems on earth have been markedly influenced by man. More than 10% of the land surface is used in agriculture. The world forests, covering about 30% of the land surface, are rapidly being exploited and grassland, savannas, and shrubland are changed by increasing herds of cattle. It becomes, therefore, important to distinguish between natural ecosystems and derived ecosystems. The most extreme of the latter are the urban systems.

A most characteristic feature of the ecosystems on land is their mosaic structure, primarily as a result of basic soil features and the distribution of climatic zones. We shall, therefore, have to distinguish between many different subsystems or biomes even though, in an attempt to model the global carbon cycle, we shall have to restrict ourselves to a rather limited number. Each one of these will, therefore, still be quite inhomogeneous and will consist of regions with different characteristics. The definition of these subsystems, the description of their dynamics, and the reduction to a few variables becomes, therefore, a fundamental task. It must precede the construction of simple dynamic ecosystem models to be incorporated into a global model of the carbon cycle.

The biogeocenosis, a basic ecologic unit of which all more aggregated systems are composed, consists of a plant association (B_B) (phyto-cenosis), on which depends the existence of animals, bacteria and fungi. Together, they form a trophic network. The green photosynthetic plants are the producers of organic material. The total quantity of carbon fixed by photosynthesis is the gross primary production, P_B. Part of this is used as an energy source for maintenance and is called respiration, R_A. The organic matter stored in plant tissues in excess of respiration is the net primary productivity, P_N (or NPP). Under very favourable conditions, R_A may be merely 10% of P_B, but on the average R_A equals approximately P_N.

P_N serves as food for the animals, the consumers, whereby a predation trophic chain is initiated. Dead organic matter from producers and consumers (litter fall) initiates a decomposition trophic chain on, and particularly in, the soil, due to the activity of small animals and microorganisms. Even though their total mass is small, they constitute an important link in the carbon cycle because of their rapid turnover. Some parts of the dead organic matter are, however, decomposed rather slowly, particularly in cold climates, and form humus.

When concerned with the global carbon cycle, we particularly need to estimate the total living biomass and its distribution over various subsystems. Also important, is the characteristic transit time distribution function (Bolin and Rodhe, 1973) of the flux of carbon molecules through the biome under consideration, i.e. what are the relative amounts produced as short-lived structures such as grass and leaves, and what is stored in wood for longer periods. Similar estimates are needed for the soil system where the transit time distribution function is of prime concern, as it gives the probability distribution of the time for the carbon incorporation as dead organic material in the soil to oxidation and release to the atmosphere. Some general idea can be obtained, in this regard, with the aid of ¹⁴C measurements of organic soil compounds, even though, usually, only the `mean transit time' or `turnover time' is obtained in this way (see Baes et al., 1976, and also the following paragraphs).

B. Biomass and Net Primary Production

Estimates of the amount of carbon in the form of living organic matter (biota) are still quite uncertain. One such estimate is summarized in Table 1.2 and others have been presented by Bazilevich et al. (1970) and Whittacker and Likens (1975). The same is true for estimates of net primary production and the decay rates of living biota. There are several reasons for these uncertainties. The definitions of the different biomes are by no means unique, and in estimating their geographic extension there is always room for subjective judgement. It is also clear that the estimates for separate ecosystem types are more uncertain than those for the living phytomass as a whole.

There is a principal difference between the estimates by Bazilevich et al. (1970) and those by other authors. The former estimate the potential biological resources, without regard to the interference by man. Their estimates, particularly those of phytomass, are, therefore, overestimates.

Furthermore, a distinction can be made between living and dead phytomass. The latter is composed of litter and standing dead wood, which, in the case of virgin forests, may be appreciable. This complicates a more close comparison between the various estimates. Presumably, standing dead wood is included in the figures given by Bazilevich et al. (1970), but probably not by Whittacker and Likens (1975) and not in Table 1.2. In view of these comments, we shall adopt a figure of 700 x 10¹⁵ g C for living phytomass and about 150 x 10¹⁵ g C for dead phytomass, but recognize that the uncertainty of these figures is at least ±10%, and possibly as much as ±20%.

We further note that the main reason for the discrepancy between the total estimates is the different values given for tropical rain forests and subtropical forests. This becomes quite critical in the subsequent discussion because the greatest changes of land biota at present take place in tropical rain forests, particularly in South America. There is no possibility of reconciling these figures at present. However, we shall use them as a starting point for the discussion of ongoing changes, as well as the basis for the inclusion of land biota into the overall model for the carbon cycle. About 85% of all land biota is contained in forests of different kinds, regardless of which estimate is used.

The figures for the net primary production of land biota vary between 53 x 10¹⁵ g C/year (Whittaker and Likens, 1975) and 78 x 10¹⁵ g C/year (Bazilevich et al., 1970) with Table 1.2 giving a value of 63 x 10¹⁵ g C/year. The differences between the estimates for different subsystems are considerably greater. Again, there is no way of judging whether one value is more accurate than another. We shall have to explore what this uncertainty implies for the overall balances in the carbon cycle.

An estimate has also been made of the annual litter fall, yielding a value of about 80% of the NPP. This value is of importance for the modelling of the dynamics of the land biota subsystem. Merely about 20% would, therefore, be accumulated each year as wood. Comparing this figure (1215 x 10¹⁵ g C/year) with the total phytomass in forests (500 x 10¹⁵ g C), which is mostly in the form of wood, we calculate the average turnover time for wood, often called 'long-lived land biota', at 3545 years.

The phytomass values quoted above are representative for the last one or two decades. Part of the discrepancies between the various estimates may be due to different assumptions about the time or time period to which it is supposed to apply. Clearly, considerable changes are occurring due to man's increasing interference with the natural ecosystems.

Chapter 6 gives a detailed account of the way in which man is modifying land biota. Chapter 7 summarizes the estimates of what the integrated effects may have been since the middle of the last century and what the present rate might be. The surveys by Bolin (1977a), Woodwell et al. (1978), and Revelle and Munk (1977) are all still incomplete in one way or another. The scatter in the figures for changes of land biota due to human intervention is even greater than that for the inventory and NPP estimates as summarized above. The likely change since the middle of the last century seems to be 3060 x 10¹⁵ g C and the annual rate, at present, is between 1 and 6 x 10¹⁵ g C. Woodwell et al. (1978) consider even higher values possible. We shall later have to explore the consequences of these ranges.

Table 1.2 Phytomass and net primary production (mean values) of the biosphere. (Duvigneaud)

					Phytomass				Net primary production
Ecosystem type			Area  106 km2		Organic matter kg/m2		Total amount Carbon (BB)  1015 g		Organic matter kg/m2 year		Total amount carbon (PN)  1015 g/year
Equatorial rain forest			11	.0	35		175		2	.2	11	.0
Tropical seasonal forest			5	.0	18		41		1	.4	3	.2
Savannas
	derived
		grass	4	.5	3		6		2	.0	4	.1
		low trees	4	.5	7		14		2	.2	4	.5
	spiny
		shrub and tree savanna	8	.0	5		18		1	.3	4	.7
		spiny forest	2	.0	11		10		1	.2	1	.1
Warm deserts
	tropical		5	.6	0	.1	0	.3	0	.01	0	.03
	subtropical-medit		4	.2	0	.3	0	.5	0	.05	0	.1
	extreme		8	.0	0	.01	0	.03	0	.001	0	.004
Sclerophyll forest			1	.3	15		9		1	.0	0	.6
	chaparrals and macchia		2	.7	7		9		0	.8	1	.0
Semideserts (scrub and grass)
	subtrop.-medit.		6	.9	1	.3	4		0	.3	0.9
	temperate		2	.3	1	.1	1	.1	0	.2	0.2
Cold deserts
	desert grey or sandy soils		1	.0	0	.4	0	.2	0	.10	0	.05
	mountainous		1	.5	0	.7	0	.5	0	.15	0	.1
Temperate grasslands
	evergreen		2	.4	2	.0	2	.2	1	.3	1	.4
	steppes
		chernozems	3	.5	2	.0	3	.2	1	.2	1	.9
		kastanozems	2	.8	1	.4	1	.8	0	.6	0	.8
		solonetz	0	.3	1	.3	0	.2	0	.5	0	.07
Temperate forests
	evergreen		4	.0	30		55		1	.7	3	.1
	deciduous		3	.0	28		38		1	.3	1	.8
	conif. plantations		2	.0	18		16		1	.7	1	.5
Peatlands			1	.5	3	.5	2	.4	0	.4	0	.3
Mixed forest			1	.5	28		20		1	.1	0	.7
Taiga
	northern		2	.1	11		11		0	.5	0	.5
	middle		6	.0	23		63		0	.7	1	.9
	south		1	.6	26		19		0	.9	0	.6
	mountainous		5	.7	17		43		0	.6	1	.5
Tundra
	grey soils		3	.8	2	.8	5		0	.25	0	.4
	mountainous		2	.9	0	.5	0	.7	0	.08	0	.1
Polar deserts			1	.5	0	.5	0	.3	0	.05	0	.03
Mangroves			0	.3	7		1	.0	1	.5	0	.2
Swamps and marshes
	temperate		0	.6	5		1	.4	2	.5	0	.7
	tropical		1	.4	12		8		4	.0	2	.6
Halophytes			0	.5	0	.2	0	.05	0	.05	0	.01
Lakes and streams			2	.0	0	.4	0	.4	0	.02	0	.02
Cultivated lands
	temperate		6	.7	1	.2	4		1	.2	3	.6
			0	.3	3	.0	0	.4	11	.0	1	.5
	tropical		8	.5	1	.4	5		1	.4	5	.4
			0	.5	4		0	.9	1	.4	0	.3
Glaciers			13	.9
Towns			1	.0	3	.5	1	.6	0	.5	0	.2
Total biosphere			148	.8			592				62	.7

The process of deforestation is continuing at an accelerated rate. The total area of earth that can be used for agriculture is hardly more than 26 x 10¹² m², i.e. not quite twice the present agricultural area (Revelle, 1976). If the present rate of deforestation and expansion of agricultural land continues, one must expect a decline during the first half of the next century, since no further land suitable for agriculture will be available.

As was referred to in Section 1.2.1, Farmer and Baxter (1974), Freyer and Wiesberg (1974) and Stuiver (1978) have shown that the ¹³C/¹²C ratio of wood has, on average, decreased significantly since the middle of the last century. The most likely explanation for this change is that more carbon previously contained in living or dead organic matter has been returned to the atmosphere, due to man's activities, than would have been the case under natural conditions. A quantitative interpretation requires a careful analysis of the fluxes in a transitory state of the carbon cycle with the aid of models. The analyses of Stuiver (1978), Bohn (1978), and Wagener (1978) yield values for the net flux of carbon to the atmosphere of about 100 x 10¹⁵ g C, in addition to that emitted by burning fossil fuel, which is in general accord with the more direct estimates discussed previously.

All figures quoted above refer to conditions during the last century or two. We know, however, that large changes have occurred since the last glaciation, during which the size of the carbon pool of living land biota may have been merely half of what it is today.

C. Carbon in the Soil

The estimates of the total amount of organic carbon in the soil, B_S, vary even more than those for living land biota (see Chapter 5, this volume). Keeling (1973a) has summarized earlier estimates and arrives at a value of 1050 x 10¹⁵ g C, which is similar to the one given by Baes et al. (1976). Bohn (1978) and our estimates, Table 1.3, give on the other hand values which are considerably greater. The major reason for this difference is the uncertainty in the estimates of the amount of peat. Earlier summaries only include rather small quantities, while Bohn gives a value of 860 x 10¹⁵ g C and Table 1.3 a figure of 882 x 10¹⁵ g C. In the boreal zone, peat decomposes very slowly. Most of the area which constitutes the present boreal zone was covered by ice prior to about 1020 thousand years ago. As a consequence, the peats encountered here are younger in age. The high abundance of peats signals a significant net transfer of carbon to the soil.

Table 1.3 Soil organic matter (dry weight) in the biomes given in Table 1.2

There is ample evidence that the bacterial decomposition of organic soil compounds is increased if land becomes used for agriculture. Paul (1976) has reported losses between 2.5 and 6.5 kg/m² from Canadian chernozemic soils during the last 60 to 75 years. Corresponding figures for grassland and tropical lands are probably lower. If we accept Revelle's and Munk's (1977) figures that about 8.5 x 10¹² m² of land has been changed into agricultural land since the middle of the last century, the likely value for the net carbon transfer to the atmosphere during the last 100 years is 1040 x 10¹⁵ g C. Considerably larger figures have been published, the highest, 200 x 10¹⁵ g C, by Reiners (1973). This figure is, however, based on a simple model and limited data and seems, therefore, rather hypothetical. It is still, however, relevant to quote this figure as it would seem to indicate that the other estimates may be considerably too low.

The changes of the ¹³C/¹²C ratio already mentioned are due both to a reduction of the extent of forests, and to the extension of agricultural land, because of the more intense use of this land and increase of the bacterial decomposition of organic soil carbon.

D. Carbon in the Freshwater System

The flux of carbon due to the freshwater cycle is dealt with in Chapter 12. In the present context, we shall only be concerned with the net flux through the soil to the sea, maintained by the hydrological cycle. We shall adopt the following approximate values for the discharge with rivers to the sea: dissolved inorganic carbon, 0.4 x 10¹⁵ g C/year; dissolved organic carbon, 0.1 x 10¹⁵ g C/year; particulate organic carbon, 0.06 x 10¹⁵ g C/year.

E. The Dynamics of Ecosystems on Land

An ecosystem may be in a state of progression, steady state, or regression, generally as a result of changing external environmental conditions. A naked or denuded site is gradually colonized in a succession of biomes (biogeocenoses) of increasing diversity and phytomass. Generally, the succession proceeds from semidesert to grassland, shrubland and to forest. The terminal state, the climax, represents a state of maximum phytomass or a maximum that, in some senses, combines large phytomass and high primary production. In this state, there is no further increase of biomass, and a balance between fixation and decomposition, and between respiration and the energy supply for maintenance of living tissues is acquired.

For a local ecosystem, a forest fire is a catastrophe which initiates renewed progression towards a climax. Considering a sufficiently large area, within which there is an approximate statistical balance between the destruction by forest fires and the progression in other areas, the region as a whole is in a climax state.

The spatial extension of biomes is primarily determined by climate and soil characteristics. Since, to a first approximation, the climatic zones follow latitudes, the biomes more or less form zonal bands. This is already clear from Köppen's classification of climate, and is well illustrated in the north-south cross section from the Arctic Sea to the warm deserts in the southern parts of the U.S.S.R. as given by Walter (1970) and modified in Figure 1.12(a). We proceed from arctic tundra, to boreal taiga and deciduous broadleaf forest. The rapid decrease of precipitation, the increase of temperature and, associated with it, the potential evaporation towards the south gradually causes forests to fade out and there is transition of steppe, semidesert and desert. As shown in Figure 1.12(a), there is an associated

Figure 1.11 Zonobiomes (ZB) and some zonoedaphobiomes (ZE) and zonoanthropobiomes (ZA) constituting the biogeosphere (after Duvigneaud, 1972). 1. ZB1 and ZE12: Dense evergreen equatorial rain forest and dense semi-deciduous subequatiorial rain forest (Pluviisilvae). 2. ZA12: High grasslow tree humid savannas. 3. ZB6: Evergreen or semi-evergreen temperate humid forests (Laurisilvae or `Chinese forests') replaced in great parts by temperate agro-ecosystems and tree plantations. 4. ZB2: Tropical broadleaved deciduous seasonal forest (Hiemisilvae) and derived savannas. 5. ZB3 and ZE23: Tropical deciduous thorn and succulent forests or woodlands (Spinisilvae), derived savannas and tropical steppes. 6. ZB4: Warm tropical deserts and semideserts. 7. ZB9: Temperate and cold continental deserts and semideserts. 8. ZB8: Temperate continental grassy meadow-steppes and steppes (Duriprata) replaced in great parts by agro-ecosystems. 9. ZB5: Evergreen sclerophyllous forests and macchia (Durisilvae and Durifruticeta). 10. ZB7: Temperate humid broadleaved deciduous forest (Aestisilvae), replaced in great part by temperate agro-ecosystems and tree plantations. 11. ZB6: Evergreen coniferous temperate hygrophyte forest. 12. ZE1014: Mixed coniferous deciduous boreal forest. 13. ZE1014: Mountain mixed coniferous deciduous boreal forest. 14. Boreal coniferous forest (Conisilvae, Aciculisilvae, taiga). 15. ZE1416: Econtone taiga-tundra. 16. ZB11: Arctic tundras. 17. ZB12: Polar or high mountain cold semidesert change of soils and the water level. Also, the amount of litter and humus varies characteristically from north to south. In the tundra, litter forms a thick layer, while humus becomes scanty because temperature conditions are not favourable for an active pedo-fauna. In the taiga, litter is still important but humus is formed more readily and the amount increases towards the south. A maximum is reached in the transition zone from forest to steppe, where the horizon of mull types may be at a depth of 3040 m caused by decaying roots and grass. When not cultivated and ploughed, these black soils constitute important carbon reservoirs. Also, there are systematic changes from north to south of the levels of calcification, gypsification, and salinization.

Figure 1.12(a) North-to-south succession of biomes in the U.S.S.R. from tundra to desert (after Schennikow in  Walter, 1970, strongly modified)

Figure 1.12(b) North-to-south succession of biomes from pole to equator (after Bazilevich et al., 1970)

The transect in Figure 1.12(a) ends in the deserts of southern U.S.S.R. It may be continued to the equator, as shown in Figure 1.12(b). There are two more maxima, at the latitudes of evergreen temperate forests and of equatorial forests (see Bazilevich et al., 1970).

It is clear that the variations of all features of the ecosystems as shown in Figures 1.12(a) and 1.12(b) are interrelated. In order to understand their characteristics, and particularly their response to environmental changes, we need to know the times required to establish approximate climax conditions for the various subsystems. The vegetation type may be established rather quickly (in a matter of decades to centuries), while the change of soil type may require considerably more time, particularly in cold climates. The podzols of the northern boreal forest, for example, have developed since the last glaciation. Thus climatic changes obviously have a profound effect on the carbon cycle and, since the amount of carbon dioxide in the atmosphere is of importance for the climate, the reverse might also be true.

1.3 FOSSIL FUEL COMBUSTION

The burning of fossil fuels (coal, oil, and gas) has been accelerating since the middle of the last century and will probably continue to do so for quite some time into the future. Keeling (1973b) has estimated the annual input to the atmosphere, primarily by using UN statistics (Figure 1.13). The figure reveals exponential growth of fossil fuel combustion of 4.3% from 1860 to 1913 and 1945 to 1973. During the interim period, the annual increase was merely 1.2%. Figure 1.13 also shows the ratio of the observed increase of CO₂ in the atmosphere and the amount emitted by fossil fuel combustion, the `airborne fraction'. It has varied during the twenty years of CO₂ observations and on the average been 54% (Keeling and Bacastow, 1977). It should be emphasized that this value is based on the assumption that no other net transfers due to man's activities have occurred, which, as will be shown, is not a valid assumption.

Since the oil crisis in 1973 another slow-down of the expansion has occurred, however slight. It is not possible to judge whether this is a temporary situation or possibly the beginning of a slower expansion in the future. This depends on the development of alternative energy sources (e.g. Burwell, 1978) and on energy conservation measures. For a more detailed discussion of the energy problem, reference is made to a useful compilation by Perry and Landsberg (1977).

Figure 1.13 Annual input of fossil fuel CO₂ to the atmosphere from 1860 until 1974 according to data by Keeling (1973b), upper curve. The annual increase of CO₂ in the atmosphere is shown in the lower solid curve from 1958 to 1974, resulting in average airborne fraction of 54% (dashed line) (after Keeling and Bacastow, 1977. Reproduced by permission of the National Academy of Sciences, Washington, D.C.)

The total emission of carbon, in the form of carbon dioxide to the atmosphere, until 1976 was about 140 x 10¹⁵ g. The estimated available resources are given by Perry and Landsberg (1977) at about 5000 x 10¹⁵ g, of which 85% is in the form of coal and the remainder is divided between oil, oilshale, tar sand, and gas. It should be emphasized that these figures are very uncertain, because the determination of what might be available for use depends on assumptions of what can technically be recovered at reasonable costs. Much larger reserves are present in the earth's crust and some estimates indicate that as much as 70008000 x 10¹⁵ g might be recoverable. In the following, we shall use the more conservative estimate of 5000 x 10¹⁵ g.

The future use of fossil fuels is difficult to predict. Keeling and Bacastow (1977) and Revelle and Munk (1977) suggest the use of a logistic distribution as a function of time. To begin with, the combustion grows exponentially, but when the total consumption starts to become significant in comparison with the total resources, the annual percentage increase declines and becomes zero when half of the resources have been used. The annual use then decreases. With a total reservoir of fossil fuels of 5000 x 10¹⁵ g, and an initial percentage increase of the combustion of 6.53, 4.53, 2.53, and 1.53%, we obtain a distribution as function of time as given in Figure 1.14. We note that the peak consumption would be reached within about 100 years, if we follow the curve of 4.5%, which was the annual increase until the first few years of this decade.

Figure 1.14 Industrial CO₂ production for various assumed patterns of fossil fuel consumption. The figures at the different curves indicate the initial (1975) percentage increase of the annual consumption. The present (1975) annual consumption is 4.6 x 10¹⁵ g C (Keeling and Bacastow, 1977. Reproduced by permission of the National Academy of Sciences, Washington, D.C.) 

1.4 A COMPOSITE MODEL FOR THE CARBON CYCLE 

1.4.1 General Considerations

Many models of the carbon cycle have been developed. It is hardly useful to add another, unless a very specific purpose for doing so is given. At this stage of our knowledge it seems most important to formulate a model that has the following characteristics:

1. It should be simple in its basic structure, but should permit the inclusion of a more detailed treatment as our knowledge about processes is improved or more data become available for verification.

2. The simplifications of reality that are necessary for arriving at reasonably simple models, should be carefully developed from the more complete knowledge about the behaviour of the various subsystems.

3. The carbon cycle depends on other biogeochemical cycles, e.g. phosphorus and nitrogen, both in the sea and on land, and oxygen and calcium in the sea. This fact offers further possibilities for verification of the model.

4. The uncertainties of model predictions should be carefully assessed and the results should not merely be given as one plausible scenario. This requires the performance of a great number of sensitivity studies.

Figure 1.15 Model of the world oceans designed to permit the description of the basic features of the ocean circulation (see Bolin, 1977b, and Chapter 15, this volume)

We are far from being able to present a comprehensive model of this kind, but the following approach is, perhaps, a step in that direction. The presentation will, to a considerable extent, be based on the computations presented by A. Björkström (Chapter 15, this volume). We shall also summarize other attempts to analyse the global carbon cycle quantitatively, particularly those by Broecker et al. (1971), Keeling (1973a), Oeschger et al. (1975), Keeling and Bacastow (1977), Revelle and Munk (1977), Oeschger and Siegenthaler (1978) and Siegenthaler et al. (1978). The detailed structure of the model, as it will be developed below, is depicted in Figure 1.15.

The carbon cycle depends on the motions of air and water and on a number of physical, chemical and biological processes. As has already been pointed out, the atmosphere may be considered as well mixed, if the time scales we consider are longer than about one year. The following discussion will be limited to such time scales and we shall, therefore, disregard the seasonal variations or the more detailed problems of carbon dioxide transfer, associated with concentration differences within the atmosphere. The total amount of carbon present in the atmosphere will be the single variable describing the atmospheric carbon dioxide. The small amounts of methane, carbon monoxide, and other carbon-containing trace gases will not be considered in the present context, since their concentrations can be assumed to remain unchanged without serious error.

The oceans, however, mix much more slowly and need to be modelled carefully, as will be discussed in the next section.

1.4.2 Transfer by Ocean Circulation

Previous models used in analysing the carbon cycle have been simple two- or three-reservoir or diffusion models. Even though they are undoubtedly useful for a first overall assessment of the role of the oceans in the carbon cycle, they suffer from serious deficiencies. We cannot depict the real ocean circulation, which makes it impossible to determine, with the aid of observations, how accurately such models describe the transfer processes. A somewhat more realistic model has been used by Broecker et al. (1971), who considered the role of the thermocline region in some detail. Few attempts have been made to deal with several tracer elements simultaneously, in order to settle more precisely the reliability of the ocean models used. Keeling and Bolin (1967, 1968) devised a three-reservoir model for this purpose, but better spatial resolution is required in order to permit modelling that describes the major features of the ocean circulation as given in Section 1.2.2. In the following, we shall use the model proposed by Bolin (1977b) and analysed in Chapter 15.

Ocean surface waters are divided into two reservoirs, cold surface water and warm surface water, both extending to about 75 m depth, i.e. the depth of the seasonal thermocline. The division line is at about 40^° latitude. Warm surface water is that part of the oceans where a well-defined permanent thermocline exists, and cold surface water comprises the rest of the ocean surface water, which more or less effectively communicates with deeper layers of the oceans by convection. The intermediate water is modelled with the aid of two reservoirs (75540 m and 5401000 m respectively) which are in convective mutual exchange with the cold surface water. The deep sea is described with the aid of eight 500-m-deep layers underneath each other, all receiving water from the cold surface water, which, in turn, gives rise to a slow upward motion, increasing in strength as continuity requires.

The model outlined above is capable of describing both the thermocline model discussed by Broecker et al. (1971) and the two-reservoir model with a varying size of the well-mixed surface reservoir as developed by Keeling (1973a).

1.4.3 Air-Sea Exchange and the Carbonate System of the Sea

There are two important aspects of the CO₂ airsea exchange that should be included in any attempt to model the global carbon cycle quantitatively:

1. The resistance for CO₂ exchange caused by the hydration-diffusion process in the top layer of the sea.

2. The change of the carbonate equilibrium in sea-water that is induced by the CO₂ addition to the atmosphereocean system.

The former aspect is most easily modelled by assigning a proper turnover time (58 years) for CO₂ in the atmosphere (see Chapter 15, this volume). The latter requires the formulation of the carbonateborate chemical equilibrium (see Keeling 1973a, and Chapter 15, this volume). It should be observed that the chemistry of the carbonate system may well be more complex than given by Keeling (1973a). Until the fundamental reactions are known, we may simulate the system by assigning experimentally determined values for the buffer factor. It is thereby possible to assess approximately the importance of an ocean that, in this regard, behaves differently from an ideal carbonate-borate solution.

Some of the parameters required to describe the chemistry of the carbonate system are temperature dependent. The division of the ocean surface waters into one reservoir for cold water and another one for warm water will permit experiments, with the aid of which we may explore how important the CO₂ solubility variation with temperature might be.

The rate of photosynthesis in surface waters is primarily limited by the supply of nutrients, particularly nitrogen and phosphorus. This, in turn, is determined by the ocean circulation (and to a minor extent by water pollution). A proper treatment of the biological activity, therefore, implies the simultaneous treatment of the phosphorus and nitrogen cycles. This has been done in a simple three-reservoir model (Keeling and Bolin, 1967, 1968) and the model outlined here and in Chapter 15 (this volume), could, in principle, be exte