CHAPTER 4 . Sedimentary Organic Matter Preservation and Atmospheric O2 Regulation
107
millimetres to a few metres of marine sediments. The rest of the internal planet is
anoxic, just as all the Earth was prior to photosynthesis (Des Marais 1997).
Respiration, the reverse of photosynthesis, releases solar energy stored in organic
matter by reaction with molecular oxygen (Eq. 4.1 reversed). In contrast to photosynthesis, aerobic respiration can occur anywhere molecular oxygen and organic matter
coexist in the presence of life. Respiration has thus allowed heterotrophic life to pervade deeply into soil profiles, sedimentary deposits and dark subterranean waters. Deep
respiration is similarly constrained by O2 input from the lighted surface of the planet,
although the delivery mechanisms for this gas depend largely on advection (or diffusion) in air and water. Once available O2 is consumed by respiration, organic matter
degradation is carried out exclusively by single-cell organisms that often specialize
both in the biochemical substrates and oxidizing agents they process. The general
pattern for anaerobic microbial degradation is to employ electron acceptors in the
order of their maximal free energy yield (e.g. NO; > Fe 3 + > Mn02> SO~- > CO 2 ), An
attending pattern is toward decreasing flexibility in substrate utilization at lower free
energy yield, such that sulphate and carbon dioxide reducers are heavily reliant on
fermenters and other microorganisms to generate the simple substrates (e.g. acetate,
formate, oxylate) they metabolize. A consequence of these combined trends is for structurally complex, hydrolysis-resistant organic substances such as lignin and algaenans
(De Leeuw and Largeau 1993) to become increasingly difficult to metabolize under
more reducing conditions.
Because preservation of organic carbon in marine sediments is essentially the only
source for release of a molar equivalent of photosynthetically produced O 2 , it follows
that ''An organic carbon paved is an O2 saved." The net global production rate of"saved"
O2 can thus be estimated by multiplying the global average sediment accumulation
rate by the average OC content of those deposits (Berner 1989; Hedges and Keil1995).
The resulting flux is on the order of 0.15 x 10 15 g OCyr-I, which is equivalent to
-0.01 X 10 15 moles OC and O2 per year (Fig. 4.1). To maintain a steady reservoir of atmospheric O 2 (38 X 10 18 moles; Van Cappellen and Ingall1996), O2 must be taken up
at the same average rate by the weathering of continental rocks and oxidation of volcanic gases. The latter sink, however, presently accounts for only -1% of O 2 uptake by
rock weathering (Holland 1973). Among rocks, over 90% of O2 uptake is by sedimentary forms (largely shales), vs. granites and basalts (Walker 1974). Oxidation of organic
matter, ferrous iron and sulphides accounts for most of the total oxygen uptake by
sedimentary rocks (Broecker 1970). Although both climate and vegetation affect the
weathering rate of sedimentary rocks, the long-term control must be the rate of continental uplift (Walker 1974; Berner and Canfield 1989; Des Marais 1997).
This close coupling presents a major challenge, because in the absence of appreciable O 2 removal by volcanic gases, it is difficult to imagine how O 2 concentrations
might be linked to tectonics in a negative feedback loop that could explain the longterm stability of the atmospheric pool. An additional challenge is that preservation
rates of organic carbon appear to have varied appreciably over the Phanerozoic, with
an especially pronounced high 300 million years B.P. during the formation of PermoCarboniferous deposits (Berner 1987; Berner and Canfield 1989). Global uplift and
weathering rates of sedimentary rocks are also likely to have varied widely over the
last 600 million years. Any mechanism that might moderate atmospheric O2 fluctuations must therefore respond to substantially changing sources and sinks on a rela-
107
millimetres to a few metres of marine sediments. The rest of the internal planet is
anoxic, just as all the Earth was prior to photosynthesis (Des Marais 1997).
Respiration, the reverse of photosynthesis, releases solar energy stored in organic
matter by reaction with molecular oxygen (Eq. 4.1 reversed). In contrast to photosynthesis, aerobic respiration can occur anywhere molecular oxygen and organic matter
coexist in the presence of life. Respiration has thus allowed heterotrophic life to pervade deeply into soil profiles, sedimentary deposits and dark subterranean waters. Deep
respiration is similarly constrained by O2 input from the lighted surface of the planet,
although the delivery mechanisms for this gas depend largely on advection (or diffusion) in air and water. Once available O2 is consumed by respiration, organic matter
degradation is carried out exclusively by single-cell organisms that often specialize
both in the biochemical substrates and oxidizing agents they process. The general
pattern for anaerobic microbial degradation is to employ electron acceptors in the
order of their maximal free energy yield (e.g. NO; > Fe 3 + > Mn02> SO~- > CO 2 ), An
attending pattern is toward decreasing flexibility in substrate utilization at lower free
energy yield, such that sulphate and carbon dioxide reducers are heavily reliant on
fermenters and other microorganisms to generate the simple substrates (e.g. acetate,
formate, oxylate) they metabolize. A consequence of these combined trends is for structurally complex, hydrolysis-resistant organic substances such as lignin and algaenans
(De Leeuw and Largeau 1993) to become increasingly difficult to metabolize under
more reducing conditions.
Because preservation of organic carbon in marine sediments is essentially the only
source for release of a molar equivalent of photosynthetically produced O 2 , it follows
that ''An organic carbon paved is an O2 saved." The net global production rate of"saved"
O2 can thus be estimated by multiplying the global average sediment accumulation
rate by the average OC content of those deposits (Berner 1989; Hedges and Keil1995).
The resulting flux is on the order of 0.15 x 10 15 g OCyr-I, which is equivalent to
-0.01 X 10 15 moles OC and O2 per year (Fig. 4.1). To maintain a steady reservoir of atmospheric O 2 (38 X 10 18 moles; Van Cappellen and Ingall1996), O2 must be taken up
at the same average rate by the weathering of continental rocks and oxidation of volcanic gases. The latter sink, however, presently accounts for only -1% of O 2 uptake by
rock weathering (Holland 1973). Among rocks, over 90% of O2 uptake is by sedimentary forms (largely shales), vs. granites and basalts (Walker 1974). Oxidation of organic
matter, ferrous iron and sulphides accounts for most of the total oxygen uptake by
sedimentary rocks (Broecker 1970). Although both climate and vegetation affect the
weathering rate of sedimentary rocks, the long-term control must be the rate of continental uplift (Walker 1974; Berner and Canfield 1989; Des Marais 1997).
This close coupling presents a major challenge, because in the absence of appreciable O 2 removal by volcanic gases, it is difficult to imagine how O 2 concentrations
might be linked to tectonics in a negative feedback loop that could explain the longterm stability of the atmospheric pool. An additional challenge is that preservation
rates of organic carbon appear to have varied appreciably over the Phanerozoic, with
an especially pronounced high 300 million years B.P. during the formation of PermoCarboniferous deposits (Berner 1987; Berner and Canfield 1989). Global uplift and
weathering rates of sedimentary rocks are also likely to have varied widely over the
last 600 million years. Any mechanism that might moderate atmospheric O2 fluctuations must therefore respond to substantially changing sources and sinks on a rela-