Isotope Geochemistry Lecture 39: Carbon Cycle & Climate II - Long-Term Cycle & Isotopes, Study notes of Geochemistry

Download Isotope Geochemistry Lecture 39: Carbon Cycle & Climate II - Long-Term Cycle & Isotopes and more Study notes Geochemistry in PDF only on Docsity! Geol. 656 Isotope Geochemistry Lecture 39 315 THE CARBON CYCLE, ISOTOPES, AND CLIMATE II THE LONG-TERM CARBON CYCLE On geologic times scales, the carbon cycle model must be augmented by 3 reservoirs, sedimentary carbonate, sedimentary organic carbon, and the mantle, as well as fluxes between these reservoirs and the oceans and atmosphere. Such a long-term model is shown in Figure 39.1, where the anthropogenic perturbations have been removed. The most important thing to notice is that there is much more car- bon in the carbonate and sedimentary organic carbon reservoirs than in all the reservoirs in Figure 38.1 combined. However, the fluxes to and from the sedimentary reservoirs are small, so they play little role in short-term (< 1 Ma) atmospheric CO2 variations (at least in natural ones: we could properly con- sider fossil fuel burning as a flux from sedimentary organic carbon to the atmosphere). We should also point out that only a small fraction of the sedimentary organic carbon is recoverable fuel; most is pre- sent as minor amounts (typically 0.5% or less) of kerogen and other refractory organic compounds in sediments. Even greater amounts of carbon are probably stored in the mantle, though the precise amount is difficult to estimate. An order of magnitude figure might be 250-500 ppm CO2 in mantle, which implies a total inventory of 2.5-5 × 108 Gt, or about 106 times the amount in the atmosphere. Again, the flux from the mantle to the atmosphere, which results from volcanism, is small, so the man- tle plays no role in short-term atmospheric CO2 variations. On long time scales (>106 yr), however, it is the fluxes to and from sediments and the mantle that control the atmospheric CO2 concentration. ATMOSPHERE 0.725 !13C –7.5 OCEAN 38.5 !13C +0.6‰ BIOSPHERE + SOIL !13C -25‰ 2 SEDIMENTARY ORGANIC CARBON !13C -20‰ 16 x 103 0.01 LongTerm Carbon Cycle Subduction 4 X 10-5 (Ca,Mg)CO3 !13C ~ +1 5.6 x 104 Volcanism 4 X 10-5 2 x 10-4 4 x 10 -4 Weathering, Sedimentation Weathering, Metamorphism 2 x 10-4 4 x 10-4 CARBONATE SEDIMENTS 0.09 0.01 0.09 !13C ~–6‰ MANTLE ~5 x 105 !13C ~–6‰ Figure 39.1. The Carbon Cycle. Green numbers show the amount of carbon (in 1018 grams) in the res- ervoirs. Fluxes between these reservoirs (arrows) are shown in italics in units of 1018 g/yr (in red). Masses and fluxes refer to the pre-Industrial Revolution state of the system. Uncertainties on many of the masses and fluxes are large. Also shown are estimates of the carbon isotopic composition. Docsity.com Geol. 656 Isotope Geochemistry Lecture 39 316 TERTIARY CARBON ISOTOPE RATIOS AND EXTINCTIONS Figure 39.2 shows δ13C and δ18O in benthic forams from 40 DSDP and ODP drill cores selected to rep- resent, as best as possible, global means. On these time scales, the main influences δ13C are changes in biological productivity and ocean circulation, burial and erosion of carbon in sediments, and the vol- canic flux. Recall that organic carbon has strongly negative δ13C – burial of organic carbon will drive the marine system toward more positive values, erosion of organic carbon will drive it to negative val- ues. Volcanic carbon has δ13C of -6, and consequently volcanism will drive the system to more negative values. The fractionation between dissolved carbonate and precipitated carbonate is fairly small, so Figure 39.2. δ13C and δ18O in marine benthic foraminifera. Data are a compilation from many cores. In the Late Miocene, δ13C of Pacific and Atlantic bottom watere diverge, and these are shown as separate curves. Also shown are significant climatic, tectonic, and biologic events Modified from Zachos et al. (2001). Docsity.com Geol. 656 Isotope Geochemistry Lecture 39 319 bon from organic or carbonate sediment through metamorphism or magmatism and weathering equals the rate burial of organic and carbonate sediment. The isotopic composition of the oceans and atmos- phere depends on these fluxes: δoFbc + (δo – αc)Fbg = δo(Fwc + Fmc) + δg(Fwg + Fmg) 39.2 where the subscript o denotes the ocean and αc is the fractionation during photosynthesis. Because the isotopic composition of the oceans through time can be estimated from δ13C in carbonate (e.g., Figure 39.3), equation 39.2 provides a constraint on these fluxes. Berner assumed that the rate of weathering of carbonate at any time depends on the ratio of land area to ocean area (ƒA(t)), biological activity (ƒE(t)), a rate constant kwc, river runoff (ƒD(t)), the mass of car- bonate rock (C), and the CO2 “weathering feedback function”, ƒCO2(t): Fwc = ƒCO2(t)ƒA(t)ƒD(t)ƒE(t)kwcC 39.3 The “weathering feedback function” works in two ways. First of all, global surface temperatures should correlated with atmospheric CO2 concentrations. Since weathering reaction rates are, in principle, tem- perature dependent, Berner reasoned that weathering would be more rapid when temperatures, and hence atmospheric CO2 concentrations, are higher (these same assumptions are present in the BLAG model; Berner et al., 1983). Second, Berner assumes that higher atmospheric CO2 leads to greater rates of photosynthesis and biological activity. This enhances weathering through greater production of bio- logical acids and nutrient uptake. One might also speculate that atmospheric CO2 might directly speed weathering since protons generated by dissociation of carbonic acid plays a key role in weathering. However, the dissolved CO2 in groundwater comes primarily from respiration by soil organisms rather than the atmosphere. Hence increasing atmospheric CO2 would not directly effect weathering rates. Berner assumed the dependence of the weathering feedback function is related to at- mospheric CO2 by the following equation of Volk (1987): ƒ CO 2 = 2R CO 2 1+ R CO 2 ! " # # $ % & & 0.4 ' 0.4 ' R CO 2 0.22 39.4 where RCO@ is the ratio of atmospheric CO2 concentration at the time of interest to the pre- sent concentration. The first term reflects the dependence of photosynthesis rate on CO2 concentration of the form seen in Figure 35.8. The 0.4 exponent represents an acknowledge- ment that biological productivity is often lim- ited by factors other than CO2 availability. The second term reflects the temperature feedback as first formulated by Berner et al. (1983). For the period before the emergence of land plants (before 350 Ma), the first term is not present. Other evolutionary changes were accounted for in the biological activity function (ƒE(t)). The rate of organic sediment weathering is given by: Fwc = ƒA(t)ƒD(t)ƒR(t)kwgG 39.5 where ƒR is a factor that depends on mean land elevation, kwg is a rate constant, and G is the mass of buried organic carbon. The terms for biological activity (ƒE(t)) ƒCO2(t) do not occur because weathering of sedimentary organic matter results simply from oxidation rather than attack by carbonic acid or bio- a Ocean and atmosphere C Carbonate C OrganicC 5000 1250 2.9 Fwc Fwg FbgFbcFmc Fmg Figure 39.5. Simple model of carbon flow consid- ered by Berner (1991). Masses of carbon are given in units of 1018 moles. Fluxes are described in the text. After Berner (1991). Docsity.com Geol. 656 Isotope Geochemistry Lecture 39 320 logically produced acids. On the other hand ƒR(t) is omitted from the expression for weathering of car- bonate because their weathering shows little dependence on elevation. Berner assumes that carbon is also deeply recycled through subduction of oceanic sediment. The metamorphic or magmatic release of CO2 from carbonate rock depends on the ratio (ƒC(t)) of platform carbonate to deep-sea carbonate (the latter being more commonly subducted) and the volcanism rate (ƒG), so: Fmc = ƒG(t)ƒC(t)kmcC 39.6 where kmc is a rate constant. Metamorphic or magmatic release of organic carbon is expressed as: Fmg = ƒG(t)kmgG 39.7 Finally, the flux of carbon due to weathering of silicate rocks and consequent uptake of CO2 and burial as carbonate (Fws) is expressed as: Fws = Fbc –Fwc = ƒCO2(t)ƒA(t)ƒD(t)ƒR(t)ƒE(t)Fws(0) 39.8 where Fws(0) is the present flux. Berner estimated the values of the various ƒ and k parameters in these equations, as well as initial (at 570 Ma) values for the sizes and isotopic composition of the three reservoirs from information in the geological literature. For example, he assumed the volcanism rate was proportional to the rate of sea- floor spreading. Factors such as mean elevation and the ratio of land to ocean area are can be estimated from geologic information. He then calculated the magmatic and weathering fluxes, and substituting these into equations 39.1 and 39.2, calculated the burial fluxes in 1 million year steps. From values of Fwc and Fbc, he solved for ƒCO2(t) in equation 39.8 and then for CO2(t). This new value of ƒCO2(t) was then used to iterate the calculation until a constant ƒCO2(t) was obtained. From this, new values for the mass of the reservoirs and their isotopic composition were calculated using the following mass balance equations such as: dC/dt = Fbg – (Fwc + Fmc) 39.9 and d(δcC)/dt = δoFbc – δc(Fwm + Fmc) 39.10 The results, with an error envelope based on the sensitivity of the method to various uncertainties in the input parameters, are shown in Figure 39.6. The results correspond more or less with what is known from the geologic record about temperature changes during the Phanerozoic. To begin with, the Early Paleozoic was warm compared with the late Precambrian, which was a time of several major glaciations. The late Paleozoic, on the other hand, was cool, and the time of the last major glacial epoch before the late Tertiary/Quaternary glaciation. The Cretaceous, on hand, is well known as a remarka- bly warm period. Berner’s model shows generally high CO2 during warm periods of the early Paleo- zoic, low CO2 (resulting from organic carbon burial, presumably a consequence of colonization of land by plants) associated with glaciation in the late Paleozoic, and high CO2 (associated with volcanism, among other things) in the warm Cretaceous. Thus if the model is correct, it substantiates the widely held assumption that atmospheric CO2 concentrations strongly influence global temperature. This is simply a model, however, and many question its validity. In particular, John Edmond has ar- gued that although weathering reaction rates, like all reaction rates, are temperature dependent, this dependence is not important in nature because other factors limit reaction rates. He argues that the most important factor limiting weathering is the abundance of fresh rock, which is in turn controlled by tectonism. He points to the Orinoco drainage as an example (Edmond et al., 1995). Although tempera- tures are high, weathering is slow because a deep layer of thoroughly weathered soil inhibits water from reaching fresh rock. While Berner’s results are certainly interesting, just how accurate these esti- mates are remains to be seen. Docsity.com Geol. 656 Isotope Geochemistry Lecture 39 321 Figure 39.6. Ratio of atmospheric CO2 concentration to present atmospheric CO2 in the model of Berner (1991). Docsity.com