10 Influence of Geochemical Processes on Stable Isotope Distribution in Marine Sediments
350
Values for δ 13 C d are estimated on the basis of the
δ 13 C of calcite tests assuming that the calcite was
precipitated in equilibrium with the ΣCO 2 , correcting for the temperature-dependent fractionation
ε b(a) between CO 2(aq) and dissolved bicarbonate
according to Mook et al. (1974):
δ 13 C d = δ 13 C ΣCO2 - ε b(a)
= δ 13 C ΣCO2 - 24.12 - 9866 / T
(10.12)
T is the temperature of surface waters measured in K. The CO 2(aq) concentrations are then
converted to CO 2 partial pressure values (pCO 2 in
µatm) using Henry’s Law:
pCO 2 = CO 2(aq) / α
(10.13)
where the solubility coefficient α is mainly a function of temperature and, to a minor extent, of
salinity (Rau et al. 1991). α may be calculated according to Weiss (1974). The success of the equations in determining surface CO 2(aq) is dependent
on the estimation of the sea-surface temperatures
and on the determination of the complex variable b
for the past. While several proxies exist for reconstructing past SST, the determination of ancient bvalues is difficult. A promising approach is the use
of bulk sediment δ 15 N as proxy for b (Andersen et
al. 1999). However, the use of δ 15 N as a proxy for
carbon demand introduces a number of potential
problems associated with the complex fractionation
mechanisms of nitrogen isotopes (see Sect. 10.5).
A topic of much concern regarding the use of
bulk δ 13 C org in paleoceanographic investigations
is the masking of the marine δ 13 C org signal by terrestrially-derived organic matter. Since terrigenous material is usually substantially depleted in 13 C (estimated mean value of –27‰, since 90% of land plants
are C 3 plants) compared to marine-derived matter
(mean value of –19‰), the carbon isotopic composition of ocean sediments have been used to trace the
origin of sedimentary organic carbon (Newman et al.
1973; Rühlemann et al. 1996). However, the underlying assumption regarding the mixing of a marine and
a terrestrial δ 13 C org end member is not always supported because of the possible variability of δ 13 C org
values produced in terrestrial systems (e.g. change
in the relative contribution of C 3 and C 4 plants).
Also, marine values may change by processes occurring in the water column (e.g. changing phytoplankton growth conditions, see above).
To circumvent the problem of influencing
carbon isotope signals by terrigenous organic matter contamination or by marine photosynthesizers
with varying carbon fixation pathways and cell
geometries, recent studies concentrated on
certain organic molecules (biomarker) from known
marine sources unique to a particular class of
organism. In this context, C 37 alkenones have
been shown to serve as excellent biomarkers (e.g.,
Andersen et al. 1999), because they derive
exclusively from some haptophyte algae, which
have a narrow range of cell geometries and
diameters (Bidigare et al. 1997; Popp et al. 1998),
thereby minimizing these variables as a
significant control on isotopic compositions. As
an application, Pagani et al. (1999) concluded on
the basis of ε p37:2 records that early to late
Miocene (25-9 Ma) pCO 2 levels were similar to
those recorded for Pleistocene glacial-interglacial
intervals. This pattern of low pCO 2 levels
throughout the Neogene has been confirmed
independently by Pearson and Palmer (2000)
using boron isotopes (see chapter 10.7)
Diagenesis
Early diagenesis begins in the photic zone of the
oceans, continues during the sinking of particles,
and is intense in the bioturbated surface layer of
sediments. Therefore, only a few percent of the
initially produced organic matter becomes buried
in the sediments (Berger et al. 1989). However,
despite the extensive loss of organic matter due to
remineralization, the carbon isotopic composition
of particulate organic matter appears to undergo
only little change (see review in Popp et al. 1997).
On the other hand, large and nonsystematic
differences have been observed between the
isotopic composition of total organic carbon and
single organic compounds, e.g. phytoplankton
biomarkers, in the same deposit (Fig. 10.5; Degens
et al. 1969). Therefore, the assumption that the
carbon isotopic composition of marine sedimentary organic matter directly or indirectly reflects
the carbon isotopic composition of phytoplanktonic matter must be taken cautiously (Popp et al.
1997). Furthermore, since marine organic matter is
easily digestible, whereas terrestrial organic
matter is rather resistant, such preferential
decomposition of organic matter has great impact
on the estimation of the average marine and terrestrial δ 13 C org component (de Lange et al. 1994).
The carbon isotopic composition of pore waters reflects both the decomposition of organic
matter, which releases 13 C-depleted CO 2 to the
350
Values for δ 13 C d are estimated on the basis of the
δ 13 C of calcite tests assuming that the calcite was
precipitated in equilibrium with the ΣCO 2 , correcting for the temperature-dependent fractionation
ε b(a) between CO 2(aq) and dissolved bicarbonate
according to Mook et al. (1974):
δ 13 C d = δ 13 C ΣCO2 - ε b(a)
= δ 13 C ΣCO2 - 24.12 - 9866 / T
(10.12)
T is the temperature of surface waters measured in K. The CO 2(aq) concentrations are then
converted to CO 2 partial pressure values (pCO 2 in
µatm) using Henry’s Law:
pCO 2 = CO 2(aq) / α
(10.13)
where the solubility coefficient α is mainly a function of temperature and, to a minor extent, of
salinity (Rau et al. 1991). α may be calculated according to Weiss (1974). The success of the equations in determining surface CO 2(aq) is dependent
on the estimation of the sea-surface temperatures
and on the determination of the complex variable b
for the past. While several proxies exist for reconstructing past SST, the determination of ancient bvalues is difficult. A promising approach is the use
of bulk sediment δ 15 N as proxy for b (Andersen et
al. 1999). However, the use of δ 15 N as a proxy for
carbon demand introduces a number of potential
problems associated with the complex fractionation
mechanisms of nitrogen isotopes (see Sect. 10.5).
A topic of much concern regarding the use of
bulk δ 13 C org in paleoceanographic investigations
is the masking of the marine δ 13 C org signal by terrestrially-derived organic matter. Since terrigenous material is usually substantially depleted in 13 C (estimated mean value of –27‰, since 90% of land plants
are C 3 plants) compared to marine-derived matter
(mean value of –19‰), the carbon isotopic composition of ocean sediments have been used to trace the
origin of sedimentary organic carbon (Newman et al.
1973; Rühlemann et al. 1996). However, the underlying assumption regarding the mixing of a marine and
a terrestrial δ 13 C org end member is not always supported because of the possible variability of δ 13 C org
values produced in terrestrial systems (e.g. change
in the relative contribution of C 3 and C 4 plants).
Also, marine values may change by processes occurring in the water column (e.g. changing phytoplankton growth conditions, see above).
To circumvent the problem of influencing
carbon isotope signals by terrigenous organic matter contamination or by marine photosynthesizers
with varying carbon fixation pathways and cell
geometries, recent studies concentrated on
certain organic molecules (biomarker) from known
marine sources unique to a particular class of
organism. In this context, C 37 alkenones have
been shown to serve as excellent biomarkers (e.g.,
Andersen et al. 1999), because they derive
exclusively from some haptophyte algae, which
have a narrow range of cell geometries and
diameters (Bidigare et al. 1997; Popp et al. 1998),
thereby minimizing these variables as a
significant control on isotopic compositions. As
an application, Pagani et al. (1999) concluded on
the basis of ε p37:2 records that early to late
Miocene (25-9 Ma) pCO 2 levels were similar to
those recorded for Pleistocene glacial-interglacial
intervals. This pattern of low pCO 2 levels
throughout the Neogene has been confirmed
independently by Pearson and Palmer (2000)
using boron isotopes (see chapter 10.7)
Diagenesis
Early diagenesis begins in the photic zone of the
oceans, continues during the sinking of particles,
and is intense in the bioturbated surface layer of
sediments. Therefore, only a few percent of the
initially produced organic matter becomes buried
in the sediments (Berger et al. 1989). However,
despite the extensive loss of organic matter due to
remineralization, the carbon isotopic composition
of particulate organic matter appears to undergo
only little change (see review in Popp et al. 1997).
On the other hand, large and nonsystematic
differences have been observed between the
isotopic composition of total organic carbon and
single organic compounds, e.g. phytoplankton
biomarkers, in the same deposit (Fig. 10.5; Degens
et al. 1969). Therefore, the assumption that the
carbon isotopic composition of marine sedimentary organic matter directly or indirectly reflects
the carbon isotopic composition of phytoplanktonic matter must be taken cautiously (Popp et al.
1997). Furthermore, since marine organic matter is
easily digestible, whereas terrestrial organic
matter is rather resistant, such preferential
decomposition of organic matter has great impact
on the estimation of the average marine and terrestrial δ 13 C org component (de Lange et al. 1994).
The carbon isotopic composition of pore waters reflects both the decomposition of organic
matter, which releases 13 C-depleted CO 2 to the