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C0011 Carbon Isotope Stratigraphy To protect the rights of the author(s) and publisher we inform you that this PDF is an uncorrected proof for internal business use only by the author(s), editor(s), reviewer(s), Elsevier and typesetter TNQ Books and Journals Pvt Ltd. It is not allowed to publish this proof online or in print. This proof copy is the copyright property of the publisher and is confidential until formal publication. Chapter 11 c0011 Carbon Isotope Stratigraphy M.R. Saltzman and E. Thomas Chapter Outline 11.1. Principles of Carbon Isotope Stratigraphy 221 11.3.3. Diagenesis 231 11.2. Spatial Heterogeneity of d13C of Dissolved 11.3.4. Global Versus Local Water Mass Signals 232 Inorganic Carbon 223 11.4. Correlation Potential and Excursions 232 11.3. Materials and Methods 224 11.5. Causes of Carbon Isotope Excursions 236 11.3.1. Depositional Setting: Deep (Pelagic) 11.6. Conclusion 237 Versus Shallow 224 Acknowledgments 237 11.3.2. Bulk Versus Component 231 References 237 s0010 11.1. PRINCIPLES OF CARBON ISOTOPE STRATIGRAPHY with the ocean-atmosphere system on longer time scales (105e106 years), contains carbon in both organic matter and p0015 The potential of marine carbonate d13C values to date and carbonate rock (limestones and dolomites) (Berner, 1990). 13 12 correlate rocks relies on the fact that their C/ C values have The carbon isotopic composition of CO2 in the atmo- p0025 varied over time, mainly as the result of partitioning of carbon sphere was about À6.4 & prior to anthropogenic fossil fuel between organic carbon and carbonate carbon reservoirs in the burning, close to the mantle value of about À6& (Sundquist lithosphere (e.g., Shackleton and Hall, 1984; Berner, 1990; and Visser, 2004) (Figure 11.1). The photosynthetic fixation Kump and Arthur, 1999; Falkowski, 2003; Sundquist and of carbon using atmospheric CO2 involves a large negative Visser, 2004). Precipitation of carbonates involves little carbon fractionation, so that all organic carbon compounds are 13 isotopic fractionation relative to dissolved inorganic carbon strongly depleted in C relative to atmospheric CO2 (DIC), and the d13C of carbonate is relatively insensitive to (e.g., Maslin and Thomas, 2003). Most Recent land plants use changes in temperature (about 0.035& per C; Lynch-Stie- the C3 photosynthetic pathway, and have d13C values glitz, 2003). Therefore the d13C of inorganically and biologi- between À23 and À33& (mean value ~ À26&). Plants in dry cally precipitated carbonate in the oceans is very close to that regions (tropical grasses, salt water grasses) use a different of the DIC in the oceans (Maslin and Swann, 2005), the largest photosynthetic pathway (C4), and have d13C values ranging reservoir in the recent ocean-atmosphere system (Figure 11.1). from À9toÀ16& (mean value ~ À13&) (Maslin and To explain changes in this isotopic signature we need to Thomas, 2003). The photosynthetic reaction pathways of consider the global carbon cycle on different time scales marine phytoplankton are less well-known; d13C values in (Falkowski, 2003; Sundquist and Visser, 2004). marine phytoplankton range between À10 and À32& (most p0020 Carbon cycling between the ocean and the atmosphere lie between À17 and À22&) depending upon temperature, occurs on time scales of <1000 years. At the present pH with values in the tropics ranging up to À13&, and at high conditions of sea water (7.5e8.3), >90% of the carbon in the latitudes as low as À32& (Sarmiento and Gruber, 2006). À 13 deep ocean is present as bicarbonate (HCO3 ). The deep The d C value of whole-ocean DIC has not been constant p0030 oceanic DIC reservoir is about 50 to 60 times as large as the over geologic time. Variations in d13C in DIC in the oceans atmospheric reservoir was in pre-industrial times over time scales of tens of thousands of years or less, as for (e.g., Ravizza and Zachos, 2003; Sundquist and Visser, instance seen in the Quaternary glacial/interglacial cycles, 2004; Sarmiento and Gruber, 2006; Houghton, 2007). can be understood in terms of redistribution of carbon among Carbon in the atmosphere is present as carbon dioxide the Earth’s surface carbon reservoirs, i.e., atmosphere, (CO2), whereas the lithosphere, which exchanges carbon oceans, biosphere and superficial sediments (see The Geologic Time Scale 2012. DOI: 10.1016/B978-0-444-59425-9.00011-1 221 Copyright Ó 2012 Elsevier B.V. All rights reserved. 10011-GRADSTEIN-9780444594259 To protect the rights of the author(s) and publisher we inform you that this PDF is an uncorrected proof for internal business use only by the author(s), editor(s), reviewer(s), Elsevier and typesetter TNQ Books and Journals Pvt Ltd. It is not allowed to publish this proof online or in print. This proof copy is the copyright property of the publisher and is confidential until formal publication. 222 The Geologic Time Scale 2012 f0010 FIGURE 11.1 The carbon reservoirs of the present (pre-industrial) carbon cycle with their carbon isotopic composition. The numbers showing the size of reservoirs are expressed in the units presently most commonly used in the literature: petagrams carbon (Pg C, 1015 g carbon). Figure after Dunkley-Jones et al., 2010. e.g., Sundquist and Visser, 2004). Over timescales of much smaller size (Sundquist and Visser, 2004; Maslin and hundreds of thousands to millions of years, variations in d13C Swann, 2005). A change in carbon isotope composition of the of DIC are mainly the results of changes in the size and rate of large oceanic DIC reservoir is thus reflected in the isotopic the exchange fluxes between the Earth’s surface carbon composition of other components of the carbon cycle, within reservoirs and the lithosphere (e.g., Berner, 1990; Kump and times on the order of circulation of the deep-sea (~1000 years) Arthur, 1999; Sundquist and Visser, 2004; Maslin and Swann, (Figure 11.1): organic matter in marine and terrestrial sedi- 2005), specifically storage in the lithosphere of varying ments (Hayes et al., 1999), plant material (e.g., Robinson and amounts of carbon as organic carbon relative to the amount Hesselbo, 2004), carbonate nodules in soils (Ekart et al., 1999), stored in carbonates. The lithospheric organic carbon reser- and carbonate in herbivore teeth (e.g., Koch et al., 1992). voir includes coal measures, oil and gas reserves, but is This coupling between carbon isotope values in DIC and p0040 dominated by dispersed organic matter (Figure 11.1). Pres- organic matter may not have been in existence in the Protero- ently the carbon out-flux from the oceans into calcium zoic, when the reservoir of dissolved and particulate organic carbonate is about 4 times as large as the out-flux of carbon carbon may have been much larger than that of DIC (Rothman into organic matter (Shackleton and Hall, 1984; Shackleton, et al., 2003, Fike et al., 2006; McFadden et al., 2008; Swanson- 1987). If relatively more/less carbon is removed from the Hysell et al., 2010), although there are potentially diagenetic oceans in organic matter (relative to carbonate), the d13C explanations for this lack of coupling (e.g., Derry et al., 2010). ½AU2 value of DIC in the whole ocean increases/decreases Until the early 1990s, changes in the sizes of the global reser- (Shackleton, 1987; Berner, 1990; Kump and Arthur, 1999; voirs were thought to have occurred on time-scales of 105 to 106 Hayes et al., 1999; Derry et al., 1992; Des Marais et al., 1992; years and more (the oceanic residence time of carbon being ½AU1 Zachos and Ravizza, 2003; Sundquist and Visser, 2004; about 105 years), because oceanic deposition and erosion and Maslin and Swann, 2005). When there is net deposition of weathering on land cannot be re-organized quickly organic matter globally, the d13C value of DIC in the whole (e.g., Shackleton, 1987; Magaritz et al., 1992; Thomas and ocean increases; when there is net oxidation of organic matter Shackleton, 1996). More recently, negative carbon isotope globally, the d13C value of DIC in the whole ocean decreases. excursions (NCIEs) have been documented with a duration of p0035 The carbon isotopic composition of DIC in the oceans is not several ten thousands to hundred thousands of years, with 13 4 only linked to the d CofCO2 in the atmosphere through transition into the NCIE possibly over <<10 years, exchange between the atmosphere and surface ocean, but also although this is still under discussion (Zachos et al., 2007, through circulation between surface and deep waters in the Cui et al., 2011). Such NCIEs include the one during the oceans: the “atmosphere is the slave of the ocean” because of its Paleocene-Eocene Thermal Maximum (PETM; ~ 55.5 Ma; 10011-GRADSTEIN-9780444594259 To protect the rights of the author(s) and publisher we inform you that this PDF is an uncorrected proof for internal business use only by the author(s), editor(s), reviewer(s), Elsevier and typesetter TNQ Books and Journals Pvt Ltd. It is not allowed to publish this proof online or in print. This proof copy is the copyright property of the publisher and is confidential until formal publication. Chapter | 11 Carbon Isotope Stratigraphy 223 e.g., Kennett and Stott, 1991; Sluijs et al., 2007), smaller NCIEs (e.g., Kroopnick, 1985; Gruber et al., 1999; Sarmiento and during other late Paleocene and early Eocene hyperthermal Gruber, 2006). This vertical gradient is due to biological events (Cramer et al., 2003; Lourens et al., 2005), and possibly activity (the “biological pump”, Raven and Falkowski, 1999). earlier events, for example the Early Cretaceous and Early Photosynthesis in the oceans is limited to the photic zone, Jurassic ~183 Ma Oceanic Anoxic Events (Jenkyns, 1985, causing depletion in 12C in DIC in the surface waters.
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