SCOPE 13 - The Global Carbon Cycle

ABSTRACT

9.1 GENERAL

Item (i) has been studied on the basis of the ¹⁴C distribution in the ocean, using different models of oceanic structure and mixing. In general, the rapid equilibration between atmophere and the mixed surface layer (between 5 and 10 years) is no matter of dispute, while the calculated deep-sea mixing rates are much more dependent on the type of model. The residence time in the deep sea is between 500 and 2000 years. The deep-sea mixing rate still seems to be a parameter which can easily be adjusted to account, at least partly, for the uptake of additional CO₂ as released from the changing biosphere (besides the fossil production). The biosphere under human impact has recently been identified as an intensive CO₂ source (Woodwell et al., 1978; Bolin, 1977). Its CO₂ production has been calculated from ¹³C measurements by Stuiver (1978), Wagener (1978), and Freyer (1978).

9.2 THE BUFFER SYSTEM

CO2 (g)CO2 (aq)	(1)
CO2 (aq) +H2O H2 CO3	(2)
H2CO3H+ + HCO3	(3)
HCO3H+ + CO23	(4)
H2O H+ + OH	(5)

 while the primary buffering is realized through the process

CO2 (aq) + H2O +CO23 2 HCO3

(6)

 Metal-carbonate ion pairing influences the effective concentrations of HCO₃ and also, to some degree, the pH (Whitfield, 1974), so that the buffering capacity also depends on the salinity of sea-water. The limitation of the buffering capacity is given by the limited supply of CO₃ (under natural conditions about 13% of the inorganic carbon present). In surface sea-water, biological activity consumes CO₂ and shifts reaction (6) to the left, so that cold surface water is depleted in HCO ₃  but rich in CO²₃, and thus supersaturated with CaCO₃. Average concentrations in cold surface water are (Broecker, 1974):

[CO²₃]= |0.00| 020 mole/l 

If this water descends (e.g. in the Weddell Sea), respiration and decay processes produce CO₂, and with increasing depth and/or increasing residence time below the euphotic zone, reaction (6) is shifted to the right. As a result, water of intermediate depth is more or less saturated with CaCO₃, while deep water is unsaturated. Average concentrations in deep water are (Pacific):

 [CO²₃] = 0.00 010 moles/l

Attack of sediments (solution of CaCO₃) can be excluded for the time scale under consideration. An accurate calculation of the buffering behaviour of sea-water (Keeling, 1973) must include the contributions of the borates, and interactions with minor species.

In global models for the prediction of future CO₂ levels, the interaction between sea-water and a variable CO₂ partial pressure is usually characterized by the ratio of fractional changes in the atmosphere and the surface water,

	P/P
	c/c

where P= CO₂ partial pressure, and c = [HCO₃] + [CO²₃] +CO₂ (aq)]. Various authors, however, use different normative values (preindustrial or actual values, respectively) which give different numerical values for that ratio. Moreover, using the same name for different definitions presents a source of confusion. Table 9.1 summarizes all definitions and names and gives numerical data.

It turns out that recent experimental data on the buffer factor is lower by (17 ± 2) % than that generally accepted in the literature (Rebello, 1978b). An evaluation of Kanwisher's (1960) measurements gives a much higher figure (see Table 9.1). The importance of the buffer factor for estimating the exchange behaviour of the ocean is obvious, and additional measurements are most desirable.

All figures of Table 9.1 depend on P and temperature. A figure which was found independent of P and temperature in the range 10 ^°C < 5 T < 30 ^°C and for 200 ppm < P 2400 ppm is

		=d ln P
	*		=3876±500
		d c

(Rebello, 1978a). * can be converted into the other definitions listed in Table 9.1: 

		PPo
		Po
	=		• CO02 • *
		ln(P/Po)
	z  =	P
			• CO02 • *
		Po
	B =	* • CO2

It is important to note the dependence of the evasion factor, , which is the figure most frequently used to describe the buffer behaviour of sea-water on P or c respectively. Figure 9.1 shows measurements on South Atlantic surface water up to 6000 ppm. The curve is an 8th-grade fitting polynome, from which the evasion factors have been calculated (Figure 9.2). The evasion factors for 25.6 ^°C are lower than those calculated by Keeling (1973) for 19.6 ^°C by about 15%. The temperature effect on , however, seems to be small, so that the buffer capacity indeed seems to be higher than presently accepted. Again, more measurements are extremely desirable.

Table 9.1 Different definitions and forms to describe the interaction between CO₂ and surface ocean water, as used by various authors. P, CO₂ = actual values; Index^° refers to preindustrial data. The given numerical values refer to conditions and data as indicated

				Numerical		Conditions for the given
Author	Definition		Name	value		numerical value
Revelle and Suess	P/P0		Revelle factor	~10		Estimated for average data of
(1957)	CO2/CO02					sea-water
Bolin and Eriksson	P/P			12.5		constant A = 2.37 mval/1
(1959)	CO2/CO2					CO2= 2.148 mmol/1
						borates neglected
Kanwisher (1960)	(His experimental results
	evaluated for the evasion factor, )			16.3		31.5‰ salinity
						10 and 20 °C (giving same result)
						330 ppm CO2
Broecker et al.	d In P		Buffer factor	9.2		average ocean surface water
(1971)	B =					330 ppm CO2
		d In CO2
Bacastow and Keeling		(Pm-Pom)/Pom	Evasion factor	9.7		330 ppm; A =constant
(1973)	=					average ocean surface water
		(CO2/CO02)/CO02
Keeling (1973)		Pm/Pom	Buffer factor	10.3		average ocean surface water
	z =					330 ppm; A =constant
		CO2/CO02
Pytkowicz (1977)			(Buffer factor)-1	10.01		A = 2.331 mval/kg sw
						19.26‰ chlorinity
						pH = 8.262; T = 17.70 °C
Rebello and Wagener		Measurements in the range 300 < p < 400 ppm		B = 7.8 ± 1.0		7 sea-water samples from different
(1976)		evaluated for the evasion factor, and the		= 8.3 ± 1.1		locations; 330 ppm
		buffer factors, z, B		z=8.87 ± 1.2

Figure 9.1 Measurements of P vs. CO_a on South Atlantic surface water sampled 1 Sept. 1977, at Barra da Tijuca near Rio de Janeiro. Total alkalinity = 2.33 mM, CO₂ = 45 ml/1; total boron = 5.0 ± 0.1 ppm. Preliminary results of Rebello (1978b). The curve is an 8th-grade fitting polynome

Figure 9.2 Evasion factors (for definition see Table 9.1) as a function of the CO₂ partial pressure. P_o = 295 ppm. Upper two curves - broken line, evaluation of Kanwisher's measurements (1960) on a sea-water sample of 31.5‰ salinity; solid line, same results corrected for 36.2‰ salinity. Keeling's data from Keeling (1973). Lower curve (Rebello, 1978b) derived from the fitting polynome in Figure 9.1

Figure 9.3 log P vs. pH as derived from equilibrium measurements of Rebello (1978b). Open circles: measurements of Figure 9.1 (South Atlantic surface water, 25.6 ^°C); full circles: North Sea water, 20 ^°C

9.3 EFFECT OF HIGHER CO₂ PARTIAL PRESSURES ON THE pH OF SEA-WATER

It has been argued that, as a consequence of higher CO₂ partial pressures, the surface layers of the ocean may become undersaturated with respect to calcium carbonate, according to equation (6), with possible catastrophic biological consequences (Bacastow and Keeling, 1973; Fairhall, 1973; Zimen and Altenhain, 1973; Broecker et al., 1971). Whitfield (1974), however, calculates that this will not happen in the foreseeable future. Any change in the carbonate/bicarbonate ratio also means a change in pH. On the assumption that at any point in time in the mixed layer the equilibrium with the atmospheric CO₂ concentration is established, and using the prediction of Machta (1972) on future CO₂ levels, Whitfield finds for the change of pH:

A.D. 1958: [CO ₂ (g)] = 313 ppm;pH = 8.24 

A.D. 2010: [CO₂ (g)] = 453.5 ppm; pH = 8.16 

This gives

	pH		0.08
		=		= 0.50
	log P		0.161

Laboratory measurements under equilibrium conditions gave for this ratio 1.0; see Figure 9.3. Both Whitfield's calculation and the experiments exclude the interaction with solid CaCO₃. From equation (7) below (for references and data see Skirrow, 1975) one finds pH/ log P = -0.88 for pH > 8 and -0.95 for pH < 8:

where CA = carbonate alkalinity = [HCO₃] + 2 [CO²₃] ; K'₁ , K'₂ = first and second apparent dissociation constant of carbonic acid; = solubility coefficient of CO₂ in sea-water.

The ecological consequences of a possible change in the calcium carbonate supersaturation, with respect to calcareous organisms, can only be examined experimentally. From general aspects, it seems unlikely that CaCO_3. supersaturation is a precondition for carbonate biomineralization for the following three reasons: (i) life processes are far out of equilibria; (ii) biomineralization depends on the active transport of ions through biomembranes driven by metabolic energy, and thus do not depend on the free energy of any carbonate reaction; (iii) calcium carbonate structures of living organisms may be covered by organic tissue, representing a barrier against fast exchange processes.

Again, it is important to note in this context the recent indications which suggest that the ocean might be a stronger sink for CO₂ than believed so far, able to take up the additional CO₂ production from the shrinking biosphere.

REFERENCES

Bacastow,R. and Keeling, C. D. (1973) Atmospheric carbon dioxide and radio-carbon in the natural carbon cycle. II: Changes from A.D. 1700 to 2070 as deduced from a geochemical model. In: Woodwell, G. M. and Pecan, E. V. (eds), Carbon in the Biosphere. AEC Symposium Series 30, 86-136. NTIS U.S. Dept. of Commerce, Springfield, Virginia.

Bolin, B. (1977) Changes of land biota and their importance for the carbon cycle. Science 196, 613-615.

Bolin, B. and Eriksson, E. (1959) Changes in the carbon dioxide content of the atmosphere and the sea due to fossil fuel combustion. In: The Atmosphere and the Sea in Motion, 130-142. Rockefeller Institute Press, New York.

Broecker, W. (1974) Chemical Oceanography, 1-214. Harcourt Brace Jovanowich, Inc., New York.

Broecker, W. S., Li, Y.-H., and Peng, T.-H. (1971) Carbon dioxide man's unseen artifact. In: Hood, D. W. (ed.), Impingement of Man on the Oceans, 287-324. Wiley Interscience, New York.

Fairhall, A. W. (1973) Accumulation of fossil CO₂ in the atmosphere and the sea. Nature 245, 20-23.

Freyer, H. D. (1978) On the ¹³ C record in tree rings. I:¹³C variations in northern hemispheric trees during the last 150 years. Preliminary evaluation of the past CO₂ increase. Unpublished manuscript. To be submitted to Tellus.

Keeling, C. D. (1973) The carbon dioxide cycle: reservoir models to depict the exchange of atmospheric carbon dioxide with the oceans and land plants. In: Rasool, S. J. (ed.), Chemistry of the Lower Atmosphere, 251-329. Plenum Press, New York, London.

Pytkowicz, R. (1977) Private communication.

Rebello, A. de Luca and Wagener, K. (1976) New measurements on the system CO₂ /sea-water. Thalassia Jugoslavica (in press).

Rebello, A. de Luca (1978b) CO₂ uptake capacity of sea-water under higher CO₂ partial pressures. (UNpublished manuscript).

Skirrow, G. (1975) The dissolved gases carbon dioxide. In: Riley, J. P. and Skirrow, G. (eds), Chemical Oceanography, vol. 2, 2nd ed., 1-192. Academic Press, London, New York, San Francisco.

Wagener, K. (1977) Total anthropogenic CO₂ production during the period 1860-1975 from carbon-13 measurements. Rad. and Environm. Biophys. (in press).

Woodwell, G. M., Whittaker, R. H., Reiners, W. A., Likens, G. E., Delwiche, C. C., and Botkin, D. B. (1978) The biota and the world carbon budget. Science 199, 141-146.