656 03Lecture37 - Geol. 656 Isotope Geochemistry Lecture 37...

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Unformatted text preview: Geol. 656 Isotope Geochemistry Lecture 37 Spring 2003 THE CARBON CYCLE, ISOTOPES, AND CLIMATE I THE SHORT-TERM CARBON CYCLE AND ANTHROPOGENIC IMPACTS There is considerable reason to believe that the Earth’s climate is linked to atmospheric CO2 concentrations. There are several lines of evidence for this. First, and perhaps, most importantly, is t h e observation that CO2 gas is transparent to visible radiation but strongly absorbs infrared r adiation. Most of the energy the Earth receives from the Sun is the visible part of the spectrum. The E arth looses an equal amount of energy in the form of infrared radiation (if it did not, the Earth’s surface would continually get hotter). Atmospheric CO2 absorbs this outgoing radiation and acts as an insulating layer, keeping the Earth warmer than it otherwise would be. In principle anyway, the h igher the atmospheric CO2 concentration, the warmer the Earth’s surface will be. This is the f amiliar greenhouse hypothesis, first proposed by Savant Ahrrenius in 1895 in a paper entitled “On the Influence of Carbonic Acid in the Air upon the Temperature of the Ground”. No one thought much about this problem until Roger Revelle and Hans Suess warned in the 1950’s that the atmospheric CO2 was increasing due to burning of fossil fuels, and that this could lead to warming of the Earth’s surface. As Figure 37.1 indicates, carbon cycles rapidly between 5 reservoirs on the surface of the Earth. O f the total carbon in these 5 reservoirs, atmospheric CO2 is only a small part. Roughly equal amounts of carbon are present in the terrestrial biosphere, the atmosphere, and the surface ocean, with somewhat more being present in soil carbon. The bulk of the surficial carbon, about 7 times as much as in the atmosphere, is dissolved in the deep ocean (mainly as H CO - ). The fluxes of carbon to and from 3 the atmosphere are large relative to the amount of CO2 in the atmosphere; indeed nearly 25% of t h e atmospheric CO2 ‘turns over’ in a year. The balance of these fluxes controls the concentration of a t mospheric CO2. The isotopic composition varies between these reservoirs, primarily due to the f ractionation during photosynthesis. In the last several hundred years, man has affected the carbon cycle through burning of fossil fuels 55 DEFORESTATION 112.5 TERRESTRIAL d13C -25‰ BIOTA 2? d13C -25‰ 560 55 Short Term Carbon Cycle 5.5 FOSSIL FUEL BURNING d13C -28‰ ATMOSPHERE 725 13C -7.5‰ d 90 40 SURFACE OCEAN 92 725 d13C +1.8‰ DEEP OCEAN 38 37600 d13C +0.6‰ BIOTA 3 4 55 LITTER, PEAT, SOIL CARBON 1400 d13C -20‰ Figure 38.1. The Carbon Cycle. Numbers in green show the amount of carbon (in 1015 grams or gigatons, Gt) in the atmosphere, oceans, terrestrial biosphere, and soil (including litter, debris, etc.). Fluxes (red) between these reservoirs (arrows) are in Gt/yr. Also shown in the approximate isotopic composition of each reservoir. Magnitudes of reservoirs and fluxes are from Schlesinger (1991), isotopic compositions are from Heimann and Maier-Reimer (1996). 277 4/29/03 Geol. 656 Isotope Geochemistry Lecture 37 Spring 2003 Figure 37.2. Atmospheric CO2 concentrations (ppm volume) measured by C. D. Keeling from 1958 to 2001 at Mauna Loa (Hawaii). Annual cycles reflect the effects of seasonal changes in photosynthesis. and clearing of forests. Both these activities can be viewed as fluxes of carbon to the atmosphere, t h e former from sedimentary organic carbon, the latter from the terrestrial biosphere. The present f ossil fuel flux is between 5 and 6 Gt per year, a reasonably well-known value, and is growing; the deforestation flux is uncertain, but 2 Gt per year is a commonly cited figure. This has resulted in a roughly 0.7% per year annual increase in the concentration of CO2 in the atmosphere (Figure37.2), as d etermined by a global system of monitoring stations, the first of which were installed by C. D. Keeling in the late 1950’s at Mauna Loa and the South Pole. This is equivalent to an increase in the mass of a t mospheric CO2 reservoir of about 3 Gt/year. This increase in atmospheric CO2 is only about 58% of the fossil fuel flux and 43% of the total estimated anthropogenic carbon flux. Thus 3 or more Gt of c arbon are “missing” in the sense they are going into some reservoir other than the atmosphere, p resumably the ocean or terrestrial biosphere. Both sources of the anthropogenic carbon flux, biospheric carbon and sedimentary organic carbon have highly negative d13C (the isotopic composition of fossil fuel burned has varied over time from d13C ≈ -24‰ in 1850 to d13C ≈ -27.3‰ in 1980 as coal has been partly replaced by oil and gas). Thus we might expect to see a decrease in the d 13C of atmospheric CO2. This is indeed observed. Based on measurements of d13C in tree rings and ice cores, the d13C of atmospheric CO2 has declined by about 1.5‰ since 1800 (e.g., Figure 37.3). This is significantly greater (up to a factor of 2 greater) than t h a t expected from burning of fossil fuel alone, which is one line of evidence that there is has been a s ignificant destruction of the terrestrial biosphere over the last 200 years. 278 4/29/03 Geol. 656 Isotope Geochemistry Lecture 37 Spring 2003 To what degree the “missing” CO2 (i.e., t h a t fraction of CO2 produced by burning fossil f uel and terrestrial biosphere destruction t hat h as not accumulated in the atmosphere) has been taken up by the oceans or by terrestrial reservoirs remains a debated question. Accurate p redictions of future increases in atmospheric CO2 require an answer, because storage of carbon in these two reservoirs is quite different. Once stored in the oceans, most carbon is unlikely to re-enter the atmosphere soon. However, i ncreases in the terrestrial biomass or detritus and soil carbon may be unique, short-lived phenomena and, furthermore, may be susceptible to continued human intervention and climate change. Several t eams o f i nvestigators h ave a t tempted use to d13C changes in the atmosphere and ocean to determine what has happened to the balance of the anthropogenic carbon. UnforFigure 37.3. Variation is d 13C in an ice core from tunately, the uncertainties involved are such Sipple Station, Antarctica (open squares; F riedli that s everal o f t hese t eams h ave a rrived a t et al., 1986) and direct atmospheric samples from somewhat different conclusions. The small conthe South Pole (crosses; Keeling et al., 1989). A fcentration gradient between hemispheres ( as ter Friedli et al. (1986). indicated by the similar CO2 concentrations a t Mauna Loa and the South Pole in Figure 37.2) requires that much of the anthropogenic CO2 be t aken up in the northern hemisphere. Based on global isotopic measurements of d13C in the atmosphere, Keeling et al. (1989) concluded that the uptake by the oceans was 2.2 Gt/year in 1980. In their model, the hemispheric gradient is explained by a large northern hemisphere oceanic sink (the North A tlantic?). Quay et al. (1992) concluded based on measurement of the depth-integrated change of d13C in the oceans from 1970 to 1990 that the oceanic uptake rate was about 2.1 Gt/year. Tans et al. (1993) used the isotopic disequilibrium between the atmosphere and surface ocean to estimate an oceanic uptake rate of less than 1 Gt/year. By comparing seasonal and latitudinal variations in atmospheric d13C, Ciais et al. (1995) concluded that the terrestrial biosphere north of 30°N took up 3.6 Gt/yr in 1992-1993, while the global ocean took up only 1.82 Gt/yr in these years. They concluded that t here was a net flux of 1.7 Gt/yr from the tropical terrestrial biosphere (30°S to 30°N) to the atmosphere in these years, presumably because of deforestation. Heimann and Maier-Reimer (1996) also used t h e rate of d13C change in the ocean to estimate an oceanic uptake rate of 2.1±0.9 Gt/yr. They also pointed out the importance of the riverine carbon flux to the ocean, which previous workers had neglected. Thus there is a range in estimates of the oceanic uptake of from 1 to 2.2 Gt/yr and a clear answer as to whether the ocean or the terrestrial biosphere is the predominate sink of the “missing” anthropogenic CO2 remains elusive. Even allowing for a generous ocean uptake of 2 Gt per year leaves at least additional 3 Gt per y ear, more than the deforestation flux, that is apparently being taken up by the terrestrial biosphere. C iais et al. (1995) concluded most of this occurs in northern hemisphere temperate and polar regions. This also consistent with the hemispheric gradient in atmospheric CO2. Since most of the fossil f uel burning occurs in the northern hemisphere, we would expect the concentration of CO2 to be s lightly higher at Mauna Loa than at the South Pole. This is indeed the case (Figure 37.2); however, t h e hemispheric gradient in less than that predicted by most models of atmospheric CO2 transport, i ndicating much of the missing CO2 must be taken up in the northern hemisphere. It would appear then that expansion of the northern hemisphere terrestrial biosphere at least b alances, and likely exceeds, deforestation, which now occurs mainly in the tropics. There are s everal possible explanations for this. These are as follows. -6 -6.2 G G -6.4 GG GG GG GG G G G -6.6 G GG D GG G GG -6.8 G d13C -7 -7.2 -7.4 D D D D D -7.6 D D D D D -7.8 -8 1700 1750 1800 1850 1900 1950 2000 Year 279 4/29/03 Geol. 656 Isotope Geochemistry Lecture 37 1. 2. 3. 4. Spring 2003 As agriculture became more efficient in the 20th century, land cleared for agriculture in Europe and North America in previous centuries has been abandoned and is returning to forest. Average global temperature has increased by over 0.5°C over the last century, perhaps as a result of rising atmospheric CO2 concentrations. This temperature increase may be producing an expansion of boreal forests. Pollution, particularly by nitrates emitted when fossil fuel is burned, may be fertilizing and enhancing growth of the biosphere. As we saw in Lecture 34, plants photosynthesize more efficiently at higher CO2 concentrations, so increasing atmospheric CO2 concentrations can, in principle, stimulate plant growth. Since most plant growth is generally limited by a vailability of nutrients such as phosphate and n itrate rather than CO2, it is unclear whether such stimulation would actually occur. However, h igher CO2 concentrations may allow plants to close their stomata somewhat. Stomata, through w hich leaves exchange gas with the atmosphere, are pathways both for CO2 into the leaf and for H 2O out of the leaf. Closing the stomata somewhat would reduce water loss and therefore may a llow plants to survive in drier climates, leading to an expansion of forests and grasslands. THE QUATERNARY CARBON ISOTOPE RECORD AND GLACIAL CYCLES In our discussion of Quaternary climate cycles, we noted the need for feedback mechanisms to amplify the Milankovitch signal and mentioned that atmospheric CO2 concentration might be one of these. Early evidence that atmospheric CO2 concentration might vary between glacial and i nterglacial epochs came from carbon isotope studies of deep-sea cores. Shackleton found that seawater d13C increased during g lacial times. He a ttributed t his isotopic -2.2 change to an increase in the t errestrial biomass that would occur as a result of, among other t hings, 0 increasing land area due to f alling sea level (there is more b iological productivity per square meter on 2.2 land than in t he ocean). T his 2.0 a Planktonic would draw down atmospheric CO2 and perhaps provide the nec1.0 essary feedback to a mplify orb bital forcing of c limate change. Benthic 0 Further evidence of varying a t mospheric CO2 came from the f irst -1.0 measurement o f CO2 concen-1 c trations in ice cores in t he l a t e 300 1970’s and early 1980’s. These data suggested atmospheric CO2 -2 200 had fallen to levels as low as 200 ppm or less during glacial epochs. 0 50 100 150 200 250 300 350 The d ata a lso suggested CO2 had risen quite rapidly at the end Age, kyr of the last glaciation. The r apid Figure 37.4. Variations in d13C in (a) planktonic f oraminifera, changes s uggested t o Broecker (b) benthic foraminifera, and the difference between planktonic (1982) that the ocean must some- and benthic foraminifera (∆d 13CP–B) in core V19-30 ( Shackleton how be involved, since it i s a and Pisias, 1985) compared with the composite d 18O record of much larger carbon reservoir and Imbrie et al. (1984). Scale on the right shows the modeled exchanges r elatively q uickly change in atmospheric CO2 concentration resulting from changing with the atmosphere. He noted biological productivity of the oceans. CO2, ppm Dd13CP–B d13CPDB, ‰ d18O , s . 280 4/29/03 Geol. 656 Isotope Geochemistry Lecture 37 Spring 2003 that one obvious mechanism, changing the solubility of CO2 in the ocean due to changing temperature (solubility of CO2 increases with decreasing temperature), would produce only about a 20 ppm d ecrease in atmospheric CO2 during glacial times, and about half this would be offset by decreasing v olume of the oceans. Broecker suggested the changes in atmospheric CO2 resulted from changing biological productivity in the oceans, in other words, the effectiveness of the biological pump. He suggested that as sea level rose, phosphorus was removed by biological processes from the ocean and d eposited on continental shelves. Because the water column is short above continental shelves, there i s less opportunity for falling organic matter to be recycled before being incorporated in the sediment. He supposed that phosphorus is the limiting nutrient in the oceans; lowering its concentration would decrease marine biological productivity and thereby allow the concentration of CO2 in the atmosphere to rise. He also suggested a test of the idea. If his hypothesis were correct, the difference in d13C between surface and deep water should decrease during interglacial epochs, since this difference is a result of 12C depletion by photosynthesis in surface waters. The test could be carried out by a nalyzing the carbon isotopic composition of benthic and planktonic forams. Shackleton a nd P isias ( 1985) carried out this test by a nalyzing Depth in Core, meters the carbon isotopic composition of 500 1000 1500 2000 a core from the Panama Basin, an 2 area with sufficiently high s edimentation rate to clearly show t h e 0 changes. They found t hat t here -2 were indeed variations in the d ifference in isotopic composition be-4 tween benthic and planktonic f oraminifera ( ∆d 13CP–B) t hat corre-6 1 lated with d 18O and therefore im-8 plied d ifferences b etween d eep and surface w ater d13C (Figure -10 37.4). These differences in turn implied d ifferences i n t he b io0 d13O ‰ logical productivity in the ocean sufficient to cause changes the a t mospheric CO2. Indeed, assuming biological a ctivity r emoves o r300 ganic carbon and carbonate in a -1 constant ratio, the differences in 280 the ∆d 13CP–B suggested changes in the concentration of atmospheric CO2 260 CO2 of over 100 ppm, similar to ppm the range observed in ice cores. 240 Shackleton a nd P isias ( 1985) 220 also performed a spectral a nalysis on the d 13C data and the ∆d 13CP-B 200 parameter. The results showed there w ere i mportant s pectral 180 components at periods of 100 kyr, 40 kyr, and 23 kyr, the now f amil0 50000 100000 150000 Age, years iar Milankovitch periods. Thus Figure 37.5. Comparison of CO2 in bubbles (gray shows a nalytid13C a nd a tmospheric C O2 a re cal uncertainties) in the Vostok ice core with temperatures calcuclearly related to climate. From lated from d D and the marine d 18O record. From Barnola et a l . this analysis, however, it is un(1987). clear what is cause and what is e f281 4/29/03 Geol. 656 Isotope Geochemistry Lecture 37 Spring 2003 fect. Looking at the phase of the d 13C variations relative to those of d18O, Shackleton and P isias found that changes in ∆d13CP–B led those of d18O, suggesting ice volume was responding to CO2 concentrations, and not visa versa. This is an interesting result, but one that is inconsistent with the m echanisms of CO2 change envisioned by both Shackleton and Broecker. The best record of late Quaternary atmospheric CO2 is that provided air bubbles preserved in t h e Vostok ice core and analyzed by Barnola et al. (1987), shown in Figure 37.5. Although the details of the record differ somewhat from that predicted by Shackleton and Pisias, the range in concentrations is rather similar to that predicted (Fig. 37.4), showing the general validity of using differences in carbon isotopes as an indicator of atmospheric CO2. In the Vostok record, an increase in CO2 a ppears to lead an increase temperature by about 1000 years when glacial epochs end. At the end of the l ast interglacial, however, CO2 appears to lag temperature by as much as 10,000 years, suggesting a complex relationship between CO2, ice volume, and climate. The exact mechanism by which atmospheric CO2 concentrations change in glacial cycles remains uncertain. As we noted above, a roughly 20ppm decrease in atmospheric CO2 concentration during g lacial times would be expected from the cooling of the oceans. The exact change will depend on how much ocean cooling occurs, however, and there is now evidence this has been underestimated (continental paleoclimatic records suggest a greater variation in temperatures in the tropics than inferred from marine carbonate records). A roughly 10 ppm increase in atmospheric CO2 should occur during glacial times because of the decrease in ocean volume. Hence these two effects should produce a net 10 ppm change, only about 10% of the change actually observed. Changes in the terrestrial biosphere, high latitude peat deposits and soil carbon, the efficiency of the oceanic biological pump, and t h e vertical circulation of the oceans, may also be important. As we saw in the previous lecture, there i s indeed evidence that the deep circulation of the ocean differs between glacial and interglacial p eriods. These changes potentially affect nutrient levels in the surface water, which in turn could a ffect the efficiency of the biological pump. Ocean circulation changes may also affect atmosphere-ocean exchange as well as the residence time of carbon in the deep ocean. Measurement of atmospheric d13C during glacial periods could help to resolve this question. Looking at Figure 37.1, we can see that since the terrestrial biosphere has lower d13C than the atmosphere, Figure 37.6. Estimation of atmospheric CO2 concentration from the observed d13Corg. Graph a shows the observed correlation between the [CO2aq] in seawater and d13Corg in modern marine p hytoplankton; b shows the dependence of [CO2aq] on partial pressure of CO2 and temperature. Using the observed correlation, the partial pressure of atmospheric CO2 can be estimated from the measured d13Corg if temperature is known and equilibration between the ocean and atmosphere is assumed. From Rau (1994). 282 4/29/03 Geol. 656 Isotope Geochemistry Lecture 37 Spring 2003 d13C IN ORGANIC CARBON AND ATMOSPHERIC CO2 CONCENTRATION -18 DDD DX DX DDX XDDX DDX X D a DDX X X XX X XX X XX DXX D X X X DXXD X D -19 XX DDDDX XXD X DX XX X X D DDX D D XX X DD DX XDD D X D DXDX X D XX D D D X X D DD D X XD D XX D X -20XX DD DX XX D D DX X XD X D D D XX D D X D X X X DX X DX D D X XXX XX X D D X XD -21 D D XD D -22 XD D -23 0 40 80 120 160 d13Corg, ‰ .. Estimated [CO2aq], µM storage of carbon in the biosphere should r aise atmospheric d13C. On the other hand, since the oceans have higher d13C than the atmosphere, transfer of carbon from the atmosphere to the ocean should lower atmospheric d13C, though the effect would be smaller. Thus f ar there are only sparse data available on t h e isotopic composition of the CO2 in ice bubbles. These measurements are difficult to make because of the limited amount of CO2 present in the ice. Leuenberger et al. (1992) found t h a t atmospheric d13C was 0.3±0.2‰ lower during the last glacial period than at present. T he data then are consistent with idea t hat t h e oceans are the principal reservoir in which CO2 is stored during glacial periods. High Latitude 16 12 8 b D XD D D DXD D D DXX D X X XD D X X DDD DD XXX XD DDXXXDXD X D X X XX DD X X DD X XXD D XD DD X D DX XX D XD X X D D DDXDDDXX XD D DDXDX X DX XX XX X XDXX X XX X X X X XXXDD X X X DX D X DX XXDD X X X XXX X DXDXX XD X X XX X D D D DX 4 0 40 Equatorial 80 120 Age, ka 160 Ice cores will be able provide information on atmospheric CO2 concentrations for at best t h e 13 last several hundred thousand years. There i s, Figure 37.7. a. Variation in d Corg in core 12392-1 however, reason to believe that climate h as from 25°N in the North Atlantic (Muller et al., undergone even more dramatic changes earlier 1983) and MD 77-169 (Fontugne and Duplessy, 1986). in Earth’s history (for example, the T ertiary b. V ariation in [CO2aq] calculated by Rau (1994) variations we considered in the previous l ec- from the data in (a) using the correlation shown in ture). What role has atmospheric CO2 concen- Figure 37.6. This is compared with the variation in trations played in these climatic v ariations? [CO2aq] calculated from the variations in CO2 in bubAn answer is important because of the need to bles in Vostok ice and estimates of surface w ater predict the climatic consequences of possible temperatures. From Rau (1994). future increases in atmospheric CO2 resulting from continued burning of fossil fuels. One possible method of determining paleo-CO2 concentrations arises from an observed r elationship between d13Corg of marine phytoplankton and the concentration of dissolved inorganic CO2 (Degens e t al., 1968; Degens, 1969). The method is illustrated in Figure 37.6: there is an observed inverse correlation between dissolved CO2 and d 13Corg of phyoplankton. We found in Lecture 28 that t h e fractionation of carbon isotopes during photosynthesis is related to CO2 concentrations (Figure 2 8.4). The reason for this, in simple terms, is that when more CO2 is a vailable, plants can afford to be more selective and therefore show a greater preference for 12C. Thus in principle at least, [CO2aq] can be estimated from measurements d13Corg. [CO2aq] in equilibrium with the atmosphere depends on t h e partial pressure of atmospheric CO2 and temperature; hence if temperature is known, the p artial pressure of CO2 can also be estimated. Figure 37.7 shows an example from Rau (1994) of the calculated [CO2aq] in surface ocean water over the last 140,000 years using the correlation in Figure 37.6. d13Crorg data are from two piston cores, one from the Indian Ocean at 10°N, the other from the Atlantic Ocean at 25°N. These are compared w ith changes in [CO2aq] predicted from the observed variations in CO2 in the Vostok core (Figure 37.5) and estimated changes in ocean surface water temperature at the equator and at high latitudes. The t otal temperature glacial-interglacial variation in equatorial surface is estimated at less than 2° C ( as noted above, this may be an underestimate), while that at high latitudes varies by about 5°C. There are, however, a number of complicating factors that may limit the usefulness of d13Corg in e stimating past variations in PCO2. For one thing, the fractionation during photosynthesis will depend 283 4/29/03 Geol. 656 Isotope Geochemistry Lecture 37 Spring 2003 on temperature. For another, t h e 0.175 procedure of Rau (1994) makes no e ffort to correct for any variations in d13C in dissolved inorganic carbon 0.150 (d 13CDIC). We expect, for example, an inverse relationship between d13CDIC J 0.125 and [CO2aq] in ocean water simply because of the effects of photosyntheE sis and respiration, and we know 0.100 E 13 d CDIC h as v aried i n t he p ast E (though t he g lacial-interglacial r/[CO 2aq] 0.075 variations a ppear t o h ave been E small). It is possible to take account 13 of variations in d CDIC by measuring 0.05 E d13C i n c arbonate f rom t he same y = 0.015x + 0.371 E sediment f raction i n w hich t h e r2= 0.971 0.025 d13Corg is measured. This was done, E for example, by Jasper and H ayes (1994). B eyond t his, d iagenetic 0.000 processes i n t he s ediment m ight 15.0 5.7 1 20.0 22.5 25.0 modify the d 13Corg, and furthermore, ∆, ‰ sedimentary inorganic carbon can be a mixture from a variety of sources Figure 37.8. Variation in the ratio of growth rate (r) to [CO2aq} and will not necessarily be represen- with the fractionation of carbon isotopes during photosynthesis. tative of that in phytoplankton. To Open symbols are laboratory experiments, closed symbol w ith avoid t his p roblem, J asper a nd error bar is range of growth rates, [CO2aq], and ∆ observed in t h e Hayes (1994) analyzed d13C in a spe- equatorial Pacific. cific organic molecule (C37 alkadienone, a lipid) known to be produced by phytoplankton. There are however, other potential problems. According to the photosynthesis model of Farquhar (1982), which we presented in Lecture 28, the isotopic fractionation during photosythesis depends on the ratio of concentration of CO2 in the atmosphere to that in the cell interior: D = a + (c i /ca) (b – a) (28.1) where a is the fractionation due to diffusion, ci is the CO2 concentration in the cell interior, ca is t h e ambient CO2 concentration, and b is the fractionation during actual photosynthetic fixation. We might expect, however, that the ratio ci/ca will depend on the photosynthetic rate: at high r ates, there will be a draw down of CO2 in the cell interior. Laws et al. (1995) suggested the interior and e xterior concentrations would be related to the photosynthetic rate (or growth rate of the cells) as: r = k1 ca – k2 ci 37.1 where r is the growth rate and k1 and k2 are two constants. Rearranging and substituting into 28.1, we have: ∆ = a + (b – a)(k1 – r/ca)/k2 37.2 Assuming the other terms are constant, this equation predicts the fractionation is proportional to r atio of the growth rate to ambient CO2 concentration. This is exactly the relationship observed by Laws et al. (1995) in experiments (Figure 37.8). Thus a determination of PCO2 from d13C in organic c arbon would appear to require a knowledge of growth rates, or the assumption that they do not v ary significantly. There appear to be other complications as well. Hinga et al. (1994) found that the fractionation during photosynthesis in culture experiments depended on pH and, furthermore, varied between species. The pH dependence may be seen in Figure 27.4 and reflects in part, the dependence of the s peciation of dissolved CO2 on pH. They found no dependence on growth rate, but the range in growth rate in 284 4/29/03 Geol. 656 Isotope Geochemistry Lecture 37 Spring 2003 their experiments was small. Consistent with these observations, Goericke and Fry (1994) observed almost no correlation between the fractionation due to photosynthesis calculated from d13C of particulate organic matter and [CO2aq] in the modern ocean. Thus it remains to be seen whether d13C in organic matter can reliably be used to estimate paleo-PCO2. In a more recent attempt at this, Arthur et al. (1998) used the difference in d13C between calcite and a specific class of organic molecules, di-unsaturated alkenones, in cores from the Pacific, Atlantic, and Indian oceans to estimate paleo-PCO2 from the latest Oligocene through the Late Miocene (25 to 8 M a). Alkenones are diagenetically resistant lipids produced only by a restricted class of marine a lgae. They found that PCO2 declined sharply at the Oligocene-Miocene boundary, which coincides with a known glacial event, and continued to decline through an episode of global warming in the mid-Miocene (about 15 Ma), reaching a low of 170 ppmv. PCO2 then recovered to 220 ppmv by 9 Ma. O verall, they found little correlation between their estimate of atmospheric PCO2 and global climate. Thus these results, if correct, suggest the role of atmospheric CO2 is far less than most paleoclimate models believe. REFERENCES AND SUGGESTIONS FOR FURTHER READING Arthur, M. A., L. Kump, M. Pagani and K. Freeman. Two problems in crust-ocean atmospheric geochemical cycles: 1) The Phanerozoic history of seawater oxygen isotope geochemistry 2) neogene global climate change and the history of PCO2. Paper presented at the GERM W orkshop , La J olla, CA, 1998. Barnola, J. M., D. Raynaud, Y. S. Korotkevich and C. Lorius, Vostok ice core provides 160,000 year r ecord of atmospheric CO2, Nature, 329, 409-414, 1987. Broecker, W. S. and G. H. Denton, The role of ocean-atmosphere reorganizations in glacial cycles, Geochim. Cosmochim. Acta, 53, 2465-2501, 1989. Broecker, W. S., Glacial to Interglacial changes in ocean chemistry, Prog. 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