Relationship Between Atmospheric Carbon Dioxide Amount and Properties of the Sea Gilbert N. Plass Department of Physics, Texas A&M University, College Station, Tex. 77843
CYPCO? = [ c o d The variation of the atmospheric COZ amount with the properties of the sea is obtained from numerical solution of the equilibrium equations. Solutions are given when: there is insufficient time for equilibrium with either CaC0, deposits or the clay sediments; there is sufficient time for equilibrium with the CaC03 deposits, but not with the clay sediments; and the time is sufficiently long for equilibrium with both CaC0, deposits and the clay sediments. The variation of the partial pressure of the atmospheric C o nis shown as a function of the total C 0 2 in the sea and atmosphere, the mean temperature of the sea, and the volume of the sea. The implications for theories of climatic change are discussed.
T
here is a continual interchange between COe in the atmosphere and the oceans. If the amount of CO? is changed in one of these systems, it must eventually change in the other. The time constant for equilibration depends on many factors such as the rate of oceanic circulation, the time required to come to equilibrium with C a C 0 8 deposits, the rate at which positive ions interact with various forms of clay, the rate at which the average temperature of the seas change, and even on the influence of the atmospheric CO, on the climate and rate of formation of glaciers. Nevertheless, equilibrium calculations of the amount of CO, in the atmosphere and oceans probably have validity over certain time scales. Calculations which neglect the CaCO, equilibria are valid for the shortest time period, perhaps a few hundred years. Over longer time periods the C a C 0 3 equilibrium must also be taken into account; under some conditions the equilibrium with various clays must be considered as well. Some calculations of the variation of the atmospheric CO, amount with the total CO, in the atmosphere-ocean system were made by Plass (1956b). More recently Bolin and Eriksson (1959) and Eriksson (1963) have made more extensive calculations. However, because of the nonlinear nature of the equations, all of their results are based on solutions valid only for first-order deviations of the quantities from their original values. Solutions are often desired for the present problem when the first-order deviations are not small; in this event their results are not valid. The purpose of this paper is to present solutions which are valid over a wide range of the parameters and thus can be used to study longterm climatic changes on the earth. Average values are used for the oceanic parameters and no attempt is made to take account of the important variations of these quantities with depth and latitude. A model that allows for these variations would be much more complex and probably is not necessary for the present purpose of studying the large-scale variations which occur over long time periods. Calculation of Carbon Dioxide Equilibrium The equations which govern the COS equilibriuin are written here in the form used by Buch (1951) and Harvey (1955). The pertinent equations are: 736 Environmental Science & Technology
[HCOaI [Cod Z[CO,l
=
=
(2)
KdHCOd/[Hl
(3)
+ [HCOJ + [Cod [HCOaI + 2 [COS]
[COzl
[AI S
=
Ki[C021/[Hl
=
0.674 b Z[CO2]
+ 82.3 Pco,
(4) (5) (6)
where the following symbols have been used: Pco2, partial pressure of CO, in the atmosphere; C Y , absorption coefficient for CO, in seawater, in m M 1.-' atm-1; [CO,], sum of concn of CO, and H 2 C 0 3 in seawater, in m M ].-I; [HCO,], concn of bicarbonate ion in seawater, in m M [CO,], concn of carbonate ion in seawater, in m M l.-I; [HI, hydrogen ion concn in seawater, in M 1 . ~ 1 ; k,, first apparent dissociation coefficient of H e C 0 3 in seawater; kz, second apparent dissociation coefficient of H 2 C 0 3 in seawater; Z[C02], total CO, in seawater, in m M 1.-1; [A], carbonate alkalinity in seawater, in M equiv I.->; S, total CO? amount in ocean-atmosphere system in units of lozograms; b = VjV, where V is the volume of the oceans and V , is their present volume. The first three equations determine the chemical equilibriums, while the next two equations are the definitions of the total COY in seawater and the carbonate alkalinity. These equations apply when there is insufficient time for the system to come to equilibrium with the C a C 0 3 in the oceans. The total C o n amount in the ocean-atmosphere system can be obtained from the total COS in the seawater plus the amount in the atmosphere from Equation 6, which agrees with the data given by Eriksson (1963). Over longer time intervals the C a C 0 3 equilibrium must also be taken into account; the appropriate set of equations is Equations 1-4 plus the following:
+ 2 [COd [A,] + 2 A[Cal S = 0.674 bZ[CO?1 + 82.3 Pco, = So + 0.674 bA[Ca] [A]
=
[HC031
L
=
=
[Ca][CO,]
(7)
(8) (9)
where the following symbols have been used: [A,], carbonate alkalinity in seawater at the present time, in m M L-l; A[Ca] = [Ca] - [Ca],, change in concn of calcium ion in seawater, in mMl.-l, where [Ca], is the concn at the present time; S, total COS amount in ocean-atmosphere system in units of 1020 grams; So, total CO, amount in ocean-atmosphere system before C a C 0 3 equilibrium in units of 1020 grams, as defined by Equation 8; L , solubility product of C a C 0 3 in mM2 1.-*. Over extremely long time periods MacIntyre (1970) has stated that the clay sediments on the sea floor control the hydrogen ion concentration by absorption of these ions while releasing sodium ions. In this case it may be assumed that [HI remains constant while [A] varies. Equations 1-4 and 7-9 also describe this equilibrium. To study large variations in the atmosphere-ocean system it is essential to have accurate solutions of thesk equations.
The set of Equations 1-6 can be reduced to a single quadratic equation for the quantity [HCOd: i(0.674 b
+ 82.3 a-') 2 Kz - 0.337 bK11 [HC03I2 + S0Ki [HCO,] + Ki[A](0.337 b[Al - So)= 0
(10) Once [HC03]is obtained from the solution of this equation, all of the other variables may be obtained from simple substitutions in the remaining equations. Thus the exact equilibrium values for the complete atmosphere-ocean system may be calculated when the time intervals are relatively short (perhaps a few hundred years). For longer time intervals, C a C 0 3 equiiibrium must also be considered. There is no closed solution to the set of nonlinear equations (1-4 and 7-9). However, a computer code was written which used the half interval search method to obtain the solution of the equations accurate to eight significant figures. If equilibrium occurs with the clay sediments, the appropriate equations become linear and can be solved easily, if it is assumed that [HI is constant.
Calculated CO, Variation The variation of the partial pressure of ( 2 0 , with the total COS in the sea and atmosphere is shown in Figure 1. The dashed curves apply when C a C 0 3 reactions are neglected and represent the solutions of Equations 1-6. The solid curves apply when CaC0, equilibrium has been esi.ablished
No CaC03 Equilibrium
and represent the exact solutions of Equations 1-4 and 7-9. Results are shown for five different average temperatures for the sea, O " , 4", 12", 20°, and 28°C. The equilibrium constants in the equations for each of these temperatures were taken from the tables and data in Harvey (1955). It was assumed from Eriksson (1963) that the total CO, amount in the sea and atmosphere is 1.4175 X 10'' kg today and that [ A ] = 0.00221 with an average temperature of the sea at 4°C at the present time. The results shown here are insensitive to minor variations in the numerical values assumed for these quantities and for the equilibrium constants. The curves for C a C 0 3 equilibrium are plotted with So (the total CO, amount before C a C 0 3 equilibrium) as defined by Equation 8 as the abscissa. The connection between S and So is discussed later. When there is insufficient time for CaC03 equilibrium the partial pressure of COS varies rapidly with the total CO, amount in the sea and atmosphere from a value of 1.35 X 10-5 to 4.39 X lo-* when S = 1 and 2, respectively, and the average temperature T of the sea is 4°C. The CO, partial pressure increases as the temperature increases as shown in Figure 1. The variation of the CO, partial pressure with So is more moderate when equilibrium is established with CaC0,; the partial pressure is 5.12 X 10-6 and 1.295 X lo-, when So = 1 and 2, respectively, and T = 4°C. It is interesting that the CO, partial pressure is more sensitive to the sea temperature when there is C a C 0 , equilibrium than when there is insufficient time for this equilibrium to be established. The volume of the seas has changed appreciably during the geologic history of the earth as water has been removed and added by glaciation and other processes. The COz partial pressure as a function of the total COz in the sea and atmosphere is shown in Figure 2 for various values of b =
CaCO3 Equilibrium
No C a C g Equilibrium
lo-$$/
TOTAL coz IN SEA AND ATMOSPHERE IKG x 10-17)+.
Figure 1. Partial pressure of atmospheric COa as a function of the total CO, amount in the sea and atmosphere in units of kg X 1O-l' Curves are given for average sea temperatures of O " , 4", 12", 20", and 28 "Cand when there is insufficient time for CaC03 equilibrium and when there is CaC03 equilibrium. The total CO1 amount for the case of CaC03 equilibriumrefers to the amount present before additional CO?is deposited or dissolved from the CaC03 sediments. The present oceanic volume is assumed.
' TOTAL COz
II
IN
2 l I.4 ' I.6 ' 1.'8 .O SEA AND ATMOSPHERE (KG x IO-")+ l
2
Figure 2. Partial pressure of atmospheric COn as a function of the total C o n amount in the sea and atmosphere before C a C 0 8 equilibrium in units of kg X lo-" Curves are given for various ratios of V / V , where V is the volume of the sea, and V , is the present volume. The average temperature of the seas is assumed to be 4°C. Volume 6, Number 8, August 1972 737
. "d0
'
'
0.44
0.98
1.02 I
'
1.06
v/ v o +
'
1.10
Figure 3. Partial pressure of atmospheric COP as a function of the ratio VjV,, where V is the volume of the sea and V, is the present volume The total C o namount in the sea and atmosphere is assumed to be 1.4175 X 1017 kg. Curves are given for various average temperatures of the sea.
VjV,. When b is less than unity, water has been withdrawn from the seas and extra CO, must enter the atmosphere. The variation of the CO, partial pressure as a function of b = VjV, is shown in Figure 3. All of these curves are for S = 1.4175. The variations of the C02 partial pressure with b are larger when there is no C a C 0 3equilibrium; the variation with temperature is greater when there is C a C 0 3 equilibrium. The variation of some of the other components of the system is also shown in order to understand the C02variations better. In Figure 4 [HC03] is shown as a function of both So and b. When the C a C 0 3 equilibrium is neglected, [HC03] tends t o approach a limiting value as the total COz amount in the sea and atmosphere becomes large (So > 1.8). The variation of these quantities with both temperature and sea volume is shown in this figure. Similarly the variation of [CO,] with temperature, sea volume, and total CO, amount is shown in Figure 5. The value of [C03]tends t o become quite small when S > 1.8 and there is no CaC03 equilibrium. When there is C a C 0 3 equilibrium, the variations are relatively small compared to those when there is no C a C 0 3equilibrium. The p H of the sea is shown in Figure 6 as a function of the temperature, sea volume, and total C02 amount. The variation of the pH with b and So is less when there is CaCO8 equilibrium than without it; however, the variation with the sea temperature is greater when there is CaC03 equilibrium. All of the results when there is C a C 0 3 equilibrium have been plotted against So,the total CO, amount in the sea and atmosphere before additional carbonate ions are dissolved from the carbonate sediments or are precipitated. The total final amount of C02 in the sea and atmosphere after C a C 0 3 equilibrium, S, is shown in Figure 7 for various values of b, T , and So. Clay sediments on the sea floor are believed by MacIntyre (1970) t o maintain the pH at nearly a constant value when
I
,
,
l
i
,
,
l
~
,
C a C g Equilibrium No Coco3 Equilibrium b=l
--
-
l
'
1.8
.
]
2.0
w
0
z
g 3 2
I
910,
I.'
2
l
b '
I.
k 1 1 . b ' 2 .o1
I.
so-
Figure 4. [ H C 0 3 ]in M/I. X l o 3 for various values of So(total CO? amount in ocean-atmosphere system before C a C 0 3 equilibrium in units of lo1' kg), b (ratio of volume of sea to present volume), T (average temperature of sea), and with and without CaC03 equilibrium 738 Environmental Science & Technology
so+ Figure 5. [CO,] in M/I. X l o 4(see legend to Figure 4)
there has been an extremely long time period for equilibrium. Hydrogen ions are exchanged with sodium ions in these sediments. In general, the variations of the partial pressure of CO, with the relevant parameters are smaller when the pH is kept constant. As an example, Equations 1-4 and 7-9 were solved with this requirement. When b = 1 and T = 4OC, the partial pressure of CO, is 1.43 X lop4 and 6.90 X when So = 1 and 2, respectively. Although this variation is somewhat smaller than that shown in Figure 1 for C a C 0 3 equilibrium, considerable changes are still possible. When b = 1 and So = 1.4175, the partial pressure of C o r is 3.52 x 10-4 and 4.91 X when T = 0°C and 28"C, respectively. When T = 4°C and So = 1.4175, the partial when b = pressure of C o r is 4.56 X 10-4 and 0.302 X 0.9 and 1.1, respectively.
Theories of Climatic Change The atmospheric carbon dioxide variations can be estimated from the curves presented here as changes occur in the average temperature of the seas, the volume of the seas, and the total amount of C o nin the atmosphere and sea. Large variations in the atmospheric C 0 2amount must have some effect on the climate through its influence on the atmospheric infrared flux. Plass (1956a, b) has calculated that for a reasonable average cloud distribution over the earth that the average surface temperature should increase 2.5OC or decrease 2.7OC when the carbon dioxide amount in the atmosphere is doubled or halved. These results assume that no other factors in the atmosphere change as a result of the CO, variations. Later calculations by Kaplan (1960, 1961) and Moller (1963) obtained results for the temperature change from CO1 variations which are in essential agreement with these considering the different methods used and the complexity of the calculations. Moller's conclusions about the effect of water vapor are based on approximate equations which have been criticized by Plass (1964).
More recently, Manabe and Strickler (1964) and Manabe and Wetherald (1967) have discussed models which include for the first time some of the complex interactions that occur within the atmosphere. They have calculated from one of their models that the equilibrium temperature of the earth's surface increases 2.92OC for clear sky conditions and 2.36"C for average cloudiness conditions when the CO, content of the atmosphere is doubled. These results are remarkably close to those obtained earlier by Plass (1956a, b), when one considers both the more reliable laboratory data available to Manabe and co-workers together with the more complex model used which preserves a reasonable lapse rate in the atmosphere. Plass (1956b) suggested that a realistic theory of climatic change should consider the effect of COz-induced temperature variations on the equilibrium between the sea and the atmosphere. This is now possible on the basis of the curves given here. For example, an increase in the total CO, amount increases the partial pressure of CO,. This in turn is expected to increase the mean temperature of the earth's surface and eventually the mean temperature of the seas. This causes further increases in the atmospheric amount, since the warmer seas cannot hold as much carbonate. The sequence of events in a glacial epoch can be interpreted from Figure 2. If there is a decrease in the total C(3, in the sea and atmosphere from some external cause (such as decreased volcanic activity and deposition of carbonates or increased weathering of igneous rocks), the atmospheric CO, amount also decreases. This brings about lower temperatures according to the CO, theory; the eventual lower temperatures of the seas in turn further decrease the atmospheric amount. If these temperatures are sufficiently
1.6
g . e & i z
1.2
Coco3 Equilibrium No CaC03 Equilibrium '
0.8
bsl
So = I 4175
-1 1.6
1.8
S+
2.0
14 13
102
106
IK)
b+ 24-
2 0-
-
T: 4OC
-
h '
04.0' I
so-
Figure 6. pH of the sea (see legend to Figure 4)
14
1
'
16
'
18
%-
1
20
Figure 7. S (total C o r amount in ocean-atmosphere system after C a C 0 3 equilibrium in units of 10" kg) as a function of So (total COZ amount in ocean-atmosphere system before CaCOa equilibrium in units of lo1'kg). (For other parameters see legend to Figure 4.) Volume 6, Number 8, August 1972 739
low, appreciable glaciation begins to form on the earth’s surface; as a consequence of the volume of water frozen in the glaciers, the volume of the oceans is reduced. As the ocean volume decreases, more COS is released into the atmosphere according to Figure 2. Eventually, enough COS will accumulate in the atmosphere to warm the surface sufficiently to start melting the glaciers. This in turn increases the volume of the seas so that they can absorb excess COS from the atmosphere. The decrease in atmospheric COSamount leads to lower surface temperatures and the cycle of glaciation starts over once again. Such a cycle would probably repeat a number of times before reaching an equilibrium state because of the long time constants involved in the processes (the time for glaciers to form and melt, the turnover time of the seas, the time for the average sea temperature to change appreciably).
Discussion The equilibrium equations for the interaction of COS between the sea and the atmosphere are solved for three different assumptions : there is insufficient time for equilibrium with either the C a C 0 3deposits or the clay sediments; C a C 0 8 equilibrium occurs, but not with the clay sediments; equilibrium occurs with both the C a C 0 3 deposits and the clay sediments. Various theories of climatic change may
be tested from these results by the calculation of the partial pressure of the atmospheric COS as a function of variations in the total CO, in the sea and atmosphere, the average temperature of the sea, and the volume of the sea.
Literature Cited Bolin, B., Eriksson, E., “Changes in the carbon dioxide content of the atmosphere and sea due to fossil combustion,” “Rossby Memorial Volume,” B. Bolin, Ed., Rockefeller Institute Press, New York, N.Y., 1959. Buch, K., Haosforskningsinstitutets Skrifter, Helsingfors, no. 151, 1951. Eriksson, E., J . Geophys. Res., 68, 3871-6 (1963). Harvey, H. W., “The Chemistry and Fertility of Sea Waters,” Cambridge University Press, London, 1955. Kaplan, L. D., Tellus, 12, 204-8 (1960). Kaplan, L. D., ibid., 13, 296-300 (1961). MacIntyre, R., Sci. Amer., 223, 104-15 (1970). Manabe, S., Strickler, R. F., J . Atmos. Sci., 21, 361-85 (1964). Manabe, S., Wetherald, R. T., ibid., 24, 241-59 (1967). Moller, F., J . Geophys. Res., 68, 3866-77 (1963). Plass, G. N., Quart. J . Roy. Meteorol. Soc., 82, 310-24 (1956a). Plass, G. N., Tellus, 8, 140-54 (1956b). Plass, G . N., Geophys. Res., 69, 1663-4 (1964). Received for reuiew Nouember 16, 1971. Accepted March 6 , 1972.
Isolation of Metal-Binding Fractions from Tobacco Smoke Condensate Vincent N. Finelli, Edward E. Menden, and Harold G . Petering’ Department of Environmental Health, College of Medicine, The University of Cincinnati, Cincinnati, Ohio 45219
w Tobacco smoke condensate (TSC)from nonfilter research cigarettes was fractionated on a weak cation exchange column [carboxymethyl cellulose in the Cu(I1) form], yielding three fractions: noncomplexing substances, protonated copper-binding ligands, and nonprotonated copper-binding ligands. Analysis for copper was done by atomic absorption spectrometry and showed the amount of complexed copper in the protonated ligand fraction to be 271 f 42 pg/cigarette and in the nonprotonated ligand fraction to be 720 i 59 pg/cigarette. Several known protonated and nonprotonated ligands were also fractionated on the cation exchanger, and their behavior was compared to that of the TSC fractions. The cation exchanger was also used in the zinc, cadmium, iron(III), and lead forms to determine the binding activity of whole TSC solutions toward these metals. Results, expressed in wmol of metal/cigarette, were copper, 14.6; zinc, 12.8; cadmium, 8.3; iron, 0.5 ; and lead, 0.5.
A
previous paper (Michael et al., 1971) reported the copper-binding activity of tobacco smoke condensate (TSC) by measuring copper chelates extractable in the organic phase in a two-phase system. The health effects of cigarette smoke and the role that trace metals may have in chemical carcinogenesis (Dixon et al., 1970), atherosclerosis (Kannel, 1971) and other chronic diseases stimulated us to study the activity of TSC constituents on 740 Environmental Science & Technology
trace metals metabolism. To study the biological effects of TSC constituents on the metabolism of zinc, copper, and iron, it seemed necessary to isolate from the total tobacco tar those agents which most likely bind transition metals. Our goal was not to isolate and identify any one specific ligand, but to group in a few fractions all those TSC constituents which have metalbinding activity. Sorption of ligands from solutions by cation exchangers containing a complexing metal ion has been reported by several investigators (Stokes and Walton, 1954; Helfferich, 1961, 1962 ; Carunchio and Grassini Strazza, 1966). Preparation in situ and isolation of zero-charge complexes passing complexing anions through a cation exchanger in the metal form have also been reported (Mitrofanova et al., 1964; Muzzarelli et al., 1969). Furthermore, Walton and others have done extensive work on the fractionation of amines and other nitrogencontaining compounds using the ligand exchange chromatography technique (Cockerel1 and Walton, 1962 ; Suryaraman and Walton, 1962; Shimomura and Walton, 1968). To separate the metal-binding constituents from whole TSC, we have used a weak cation exchanger, carboxymethyl cellulose in the Cu(I1) form. This method permitted us to fractionate TSC into three distinct fractions : noncomplexing fraction (Fl)containing TSC constituents which do not bind Cu(I1); protonated ligands (FS)containing those constituents able to form zero-charge chelates with Cu(I1); nonprotonated ligands
To whom correspondence should be addressed.
