351
pore water, and the dissolution of CaCO 3 , which
adds CO 2 relatively enriched in 13 C. Since the depletion of organic matter δ 13 C exceeds by far the
range of δ 13 C exhibited in sedimentary carbonate,
the net result of these two processes is to make
dissolved CO 2 in pore water isotopically lighter
compared to the overlying bottom water. Below
the sediment/water interface, where pore water
has a δ 13 C-value near that of seawater, a δ 13 C-gradient exists in the uppermost few centimeters.
McCorkle et al. (1985) and McCorkle and Emerson
(1988) showed that the gradient of δ 13 C-profiles
observed in several box cores vary systematically
with the rain of organic matter to the seafloor. In
slowly deposited oxic sediments containing little
amounts of organic carbon, only a small decrease
in δ 13 C relative to seawater δ 13 C occurs. In
organic-carbon rich sediments, the rates of
organic matter decomposition and the production
of reduced nitrogen compounds would rapidly
decrease the pore water δ 13 C. In these anoxic
sediments, the situation is even more complex due
to bacterial methanogenesis, which follows the
sulfate reduction during early diagenesis. Since
methane-producing bacteria metabolize methane
highly enriched in 12 C (–50‰ to –100‰; Deines
1981), the pore water becomes significantly
enriched in 13 C. Such trends in δ 13 C of total
dissolved CO 2 from pore waters within anoxic
sediments have been recovered in various Deep
Sea Drilling sites (Arthur et al. 1983). However,
because pore waters are not truly closed systems,
other factors like carbon loss due to upward diffusion of methane and other dissolved carbon
species as well as precipitation of authigenic carbonates may also affect pore water δ 13 C. For example, if significant amounts of methane are
utilized in sulfate reduction, the rate of δ 13 C
decrease associated with the ΣCO 2 increase would
be anomalously high (Arthur et al. 1983).
10.4.3 δ
δ δ
δ δ 13 C in Marine Carbonates
Principles of Fractionation
Precipitation of carbonate, largely from total dissolved carbon which is present primarily as HCO 3
- ,
involves a much smaller fractionation of carbon
isotopes compared to the photosynthetic fixation
of carbon. In fact, the δ 13 C of calcite is relatively
insensitive to changes in temperature (e.g.
≈0.035‰ °C -1 ; Emrich et al. 1970) such that carbonate minerals can be used generally to monitor
changes in the δ 13 C ΣCO2 in the waters from which
they precipitate. However, the possible effects of
biologically mediated carbonate precipitation have
been considered already for oxygen isotopes (see
Sect. 10.3.2). Observations summarized by Wefer
and Berger (1991) reveal that most organisms exhibit δ 13 C-values which are in disequilibrium with
the δ 13 C ΣCO2 of the waters, from which shell carbonate is precipitated (Fig. 10.3). As discussed for
the influences on oxygen isotopic composition in
carbonate, there may be several reasons for these
deviations. McConnaughy et al. (1997) distinguish
metabolic and kinetic effects on the carbon isotopic disequilibria. Kinetic isotope effects result
from the discrimination against 13 C during hydration and hydroxylation of CO 2 . Such effects
appear to be associated with rapid calcification,
as observed in corals. Metabolic effects apparently result from changes in the δ 13 C ΣCO2 of the
microenvironment due to photosynthesis and respiration. For example, infaunal benthic foraminifera
precipitate their calcite in equilibrium with ambient
pore water δ 13 C ΣCO2 , and therefore monitor the
vertical carbon isotope gradient due to organic
matter remineralization rather than bottom water
δ 13 C ΣCO2 (Fig. 10.7; McCorkle et al. 1990, 1997).
Mackensen et al. (1993) discuss several ways, how
the δ 13 C of even epibenthic living foraminifer species could be influenced by the decay of organic
matter. An example for the consequences of such a
'Mackensen' effect is discussed by Bickert and
Wefer (1999) for the late Quaternary reconstruction of South Atlantic deep water circulation. An
extreme depletion in δ 13 C is recorded in
endobenthic foraminifers sampled from ODP Site
680. Wefer et al. (1994) attributed these excursions
to the release of methane in the continental margin
off Peru sediments, which lowers substantially the
carbon isotopic composition of CO 2 in the
porewater. Furthermore, the laboratory experiments with live planktonic foraminifera which
demonstrated symbiont photosynthetic effect on
shell δ 18 O values (Spero 1992; Spero and Lea
1993) revealed a similar effect on shell δ 13 C values.
Spero et al. (1997) obtained from experiments with
symbiotic (O. universa) and non-symbiotic (G.
bulloides) foraminifer species a δ 13 C/[CO 3
2- ]
slope of –0.007‰ µmol -1 kg -1 . Furthermore, Bemis
et al. (2000) showed a clear disequilibrium effect
related to the calcification temperature, which has
recently been approved in a field study by King
and Howard (2004). These effects might partly be
responsible for the shifts observed in foraminiferal
10.4
Geochemical Influences on 13 C/ 12 C Ratios
pore water, and the dissolution of CaCO 3 , which
adds CO 2 relatively enriched in 13 C. Since the depletion of organic matter δ 13 C exceeds by far the
range of δ 13 C exhibited in sedimentary carbonate,
the net result of these two processes is to make
dissolved CO 2 in pore water isotopically lighter
compared to the overlying bottom water. Below
the sediment/water interface, where pore water
has a δ 13 C-value near that of seawater, a δ 13 C-gradient exists in the uppermost few centimeters.
McCorkle et al. (1985) and McCorkle and Emerson
(1988) showed that the gradient of δ 13 C-profiles
observed in several box cores vary systematically
with the rain of organic matter to the seafloor. In
slowly deposited oxic sediments containing little
amounts of organic carbon, only a small decrease
in δ 13 C relative to seawater δ 13 C occurs. In
organic-carbon rich sediments, the rates of
organic matter decomposition and the production
of reduced nitrogen compounds would rapidly
decrease the pore water δ 13 C. In these anoxic
sediments, the situation is even more complex due
to bacterial methanogenesis, which follows the
sulfate reduction during early diagenesis. Since
methane-producing bacteria metabolize methane
highly enriched in 12 C (–50‰ to –100‰; Deines
1981), the pore water becomes significantly
enriched in 13 C. Such trends in δ 13 C of total
dissolved CO 2 from pore waters within anoxic
sediments have been recovered in various Deep
Sea Drilling sites (Arthur et al. 1983). However,
because pore waters are not truly closed systems,
other factors like carbon loss due to upward diffusion of methane and other dissolved carbon
species as well as precipitation of authigenic carbonates may also affect pore water δ 13 C. For example, if significant amounts of methane are
utilized in sulfate reduction, the rate of δ 13 C
decrease associated with the ΣCO 2 increase would
be anomalously high (Arthur et al. 1983).
10.4.3 δ
δ δ
δ δ 13 C in Marine Carbonates
Principles of Fractionation
Precipitation of carbonate, largely from total dissolved carbon which is present primarily as HCO 3
- ,
involves a much smaller fractionation of carbon
isotopes compared to the photosynthetic fixation
of carbon. In fact, the δ 13 C of calcite is relatively
insensitive to changes in temperature (e.g.
≈0.035‰ °C -1 ; Emrich et al. 1970) such that carbonate minerals can be used generally to monitor
changes in the δ 13 C ΣCO2 in the waters from which
they precipitate. However, the possible effects of
biologically mediated carbonate precipitation have
been considered already for oxygen isotopes (see
Sect. 10.3.2). Observations summarized by Wefer
and Berger (1991) reveal that most organisms exhibit δ 13 C-values which are in disequilibrium with
the δ 13 C ΣCO2 of the waters, from which shell carbonate is precipitated (Fig. 10.3). As discussed for
the influences on oxygen isotopic composition in
carbonate, there may be several reasons for these
deviations. McConnaughy et al. (1997) distinguish
metabolic and kinetic effects on the carbon isotopic disequilibria. Kinetic isotope effects result
from the discrimination against 13 C during hydration and hydroxylation of CO 2 . Such effects
appear to be associated with rapid calcification,
as observed in corals. Metabolic effects apparently result from changes in the δ 13 C ΣCO2 of the
microenvironment due to photosynthesis and respiration. For example, infaunal benthic foraminifera
precipitate their calcite in equilibrium with ambient
pore water δ 13 C ΣCO2 , and therefore monitor the
vertical carbon isotope gradient due to organic
matter remineralization rather than bottom water
δ 13 C ΣCO2 (Fig. 10.7; McCorkle et al. 1990, 1997).
Mackensen et al. (1993) discuss several ways, how
the δ 13 C of even epibenthic living foraminifer species could be influenced by the decay of organic
matter. An example for the consequences of such a
'Mackensen' effect is discussed by Bickert and
Wefer (1999) for the late Quaternary reconstruction of South Atlantic deep water circulation. An
extreme depletion in δ 13 C is recorded in
endobenthic foraminifers sampled from ODP Site
680. Wefer et al. (1994) attributed these excursions
to the release of methane in the continental margin
off Peru sediments, which lowers substantially the
carbon isotopic composition of CO 2 in the
porewater. Furthermore, the laboratory experiments with live planktonic foraminifera which
demonstrated symbiont photosynthetic effect on
shell δ 18 O values (Spero 1992; Spero and Lea
1993) revealed a similar effect on shell δ 13 C values.
Spero et al. (1997) obtained from experiments with
symbiotic (O. universa) and non-symbiotic (G.
bulloides) foraminifer species a δ 13 C/[CO 3
2- ]
slope of –0.007‰ µmol -1 kg -1 . Furthermore, Bemis
et al. (2000) showed a clear disequilibrium effect
related to the calcification temperature, which has
recently been approved in a field study by King
and Howard (2004). These effects might partly be
responsible for the shifts observed in foraminiferal
10.4
Geochemical Influences on 13 C/ 12 C Ratios