10 Influence of Geochemical Processes on Stable Isotope Distribution in Marine Sediments
358
is small during the entrance of sulfate into the cell (-3
to 0‰), the subsequent breaking of the S-O bonds
is causing a large change
in the isotope
composition of the reactants (SO 4
2=> SO 3
2: 22‰,
SO 3
2=> H 2 S: up to 18‰; see Canfield 2001). As a
result, sedimentary sulfide is depleted in
34
S relative
to ocean water sulfate (Kaplan and Rittenberg 1964).
The depletion is usually in the order of 20‰ to 70‰
(Hartmann and Nielsen 1969; Ohmoto 1990),
although cultural experiments yielded only depletions between 10‰ and 30‰, with a maximum reported value of 49‰ (Kaplan and Rittenberg 1964;
Bolliger 2001). The much larger fractionation
ocurring under natural conditions may be generated by a concurrent recycling of H 2 S via an intermediate oxidation to elemental sulfur or thiosulfate by
buried Fe(III) and a disproportionation of these intermediates to SO 4
2 -
and H 2 S (Jørgensen 1990;
Habicht and Canfield 1997). The disproportionation
shunts produce
34
S-enriched SO 4
2 -
and
34
S-depleted
H 2 S and thereby increase the resulting isotopic difference between SO 4
2 -
and H 2 S in the sediment pore
water (for a recent discussion see Canfield 2001).
The extent of depletion during bacterial sulfate
reduction is affected by the magnitude of reduction
rates and by the availability of sulfate. Generally, an
increasing rate of sulfate reduction leads to a
decrease in fractionation. Furthermore, if the
availability of sulfate is reduced, e.g. through
limited diffusion in non-bioturbated or impermeable
sediments, the isotopic composition of both
reactant and product shifts towards higher values.
Such a Rayleigh fractionation process has been
shown by Hartmann and Nielsen (1969) for
dissolved sulfate and sulfide in the pore water of a
core retrieved from the western Baltic Sea (Fig.
10.11). The δ
34
S-values of both reactant and
product increase with depth, because the two
isotopes in sulfate or in sulfide do not diffuse in
the same proportion as they occur in the pore
water (Jørgensen 1979).
Fractionation also occurs during processes in
the oxidative part of the sulfur cycle. Sulfide
oxidation, which is pervasive in marine environments
(for example, 90% or more of the sulfide produced
during sulfate reduction in coastal sediments are
reoxidized) includes various processes, which are all
associated with inverse isotope fractionation effects.
These processes include the phototrophic oxidation
by a variety of anoxygenic bacteria (-2 to 0‰), the
non-phototrophic oxydation (up to -18‰), and the
already mentioned disproportionation of sulfur
compounds with intermediate oxidation states as
elemental sulfur or thiosulfate (up to -33 ‰)
(Canfield 2001).
Modern Range of Values and Historical
Variability
The isotopic composition of sulfur in modern
ocean sulfate is constant within very narrow limits
and is represented by a δ 34 S-value of +20‰ with a
standard deviation of ±0.12‰ (Longinelli 1989).
The small isotopic range over a great variety of
localities and at various depths is explained with
the residence time of sulfate in the ocean which is
well in excess of the ocean mixing time (Holland
1978). This sulfate homogeneity should be considered for the evaluation of isotope data analyzed
in ancient deposits.
The sulfur isotopic composition in oceanic sulfate
through Phanerozoic time shows a Cambrian maximum
of about +30‰, a decrease to a Permian minimum of
about +10‰ and an increase towards the modern
ocean value of +20‰ (Claypool 1980; Strauss 1997).
Superimposed over this long-term secular variation is
a pronounced short-term variability exhibiting about
the same isotopic range as documented for the entire
Phanerozoic. The observed variations are assumed to
reflect changes in the sulfur redox cycle within the
ocean. Since bacterial sulfate reduction and the subsequent formation of sedimentary pyrite is associated
with a substantial fractionation, any changes in the
isotopic mass balance are generally interpreted as
changes in the burial of reduced sulfur. Furthermore,
both the rate of sulfur input and its isotopic composition may vary as a function of time depending on the
weathering rates of different sedimentary rocks, or on
the intensity of volcanic activity. For example, increased erosion of black shales containing 34 S-depleted sulfides may lower the δ 34 S value of oceanic
sulfate. On the other hand, the increased reduction of
sulfate by bacteria and removal of 34 S-depleted
sulfide from the ocean may increase its δ 34 S value.
However, the fact that the deposition of evaporites in
the Phanerozoic history occurred in rather short
episodes raises the question about the validity of the
sulfate isotope curve with respect to representativeness and homogeneity. The specific causes for the
variation of δ 34 S of marine sulfate at a particular
geologic time interval are still largely speculative (for
further discussion see Strauss 1997).
A special case in the modern ocean sulfur isotope distribution is the isotopic composition of hydrogen sulfide in anoxic basins. Fry et al. (1991)
compared depth profiles of the Black Sea and the
358
is small during the entrance of sulfate into the cell (-3
to 0‰), the subsequent breaking of the S-O bonds
is causing a large change
in the isotope
composition of the reactants (SO 4
2=> SO 3
2: 22‰,
SO 3
2=> H 2 S: up to 18‰; see Canfield 2001). As a
result, sedimentary sulfide is depleted in
34
S relative
to ocean water sulfate (Kaplan and Rittenberg 1964).
The depletion is usually in the order of 20‰ to 70‰
(Hartmann and Nielsen 1969; Ohmoto 1990),
although cultural experiments yielded only depletions between 10‰ and 30‰, with a maximum reported value of 49‰ (Kaplan and Rittenberg 1964;
Bolliger 2001). The much larger fractionation
ocurring under natural conditions may be generated by a concurrent recycling of H 2 S via an intermediate oxidation to elemental sulfur or thiosulfate by
buried Fe(III) and a disproportionation of these intermediates to SO 4
2 -
and H 2 S (Jørgensen 1990;
Habicht and Canfield 1997). The disproportionation
shunts produce
34
S-enriched SO 4
2 -
and
34
S-depleted
H 2 S and thereby increase the resulting isotopic difference between SO 4
2 -
and H 2 S in the sediment pore
water (for a recent discussion see Canfield 2001).
The extent of depletion during bacterial sulfate
reduction is affected by the magnitude of reduction
rates and by the availability of sulfate. Generally, an
increasing rate of sulfate reduction leads to a
decrease in fractionation. Furthermore, if the
availability of sulfate is reduced, e.g. through
limited diffusion in non-bioturbated or impermeable
sediments, the isotopic composition of both
reactant and product shifts towards higher values.
Such a Rayleigh fractionation process has been
shown by Hartmann and Nielsen (1969) for
dissolved sulfate and sulfide in the pore water of a
core retrieved from the western Baltic Sea (Fig.
10.11). The δ
34
S-values of both reactant and
product increase with depth, because the two
isotopes in sulfate or in sulfide do not diffuse in
the same proportion as they occur in the pore
water (Jørgensen 1979).
Fractionation also occurs during processes in
the oxidative part of the sulfur cycle. Sulfide
oxidation, which is pervasive in marine environments
(for example, 90% or more of the sulfide produced
during sulfate reduction in coastal sediments are
reoxidized) includes various processes, which are all
associated with inverse isotope fractionation effects.
These processes include the phototrophic oxidation
by a variety of anoxygenic bacteria (-2 to 0‰), the
non-phototrophic oxydation (up to -18‰), and the
already mentioned disproportionation of sulfur
compounds with intermediate oxidation states as
elemental sulfur or thiosulfate (up to -33 ‰)
(Canfield 2001).
Modern Range of Values and Historical
Variability
The isotopic composition of sulfur in modern
ocean sulfate is constant within very narrow limits
and is represented by a δ 34 S-value of +20‰ with a
standard deviation of ±0.12‰ (Longinelli 1989).
The small isotopic range over a great variety of
localities and at various depths is explained with
the residence time of sulfate in the ocean which is
well in excess of the ocean mixing time (Holland
1978). This sulfate homogeneity should be considered for the evaluation of isotope data analyzed
in ancient deposits.
The sulfur isotopic composition in oceanic sulfate
through Phanerozoic time shows a Cambrian maximum
of about +30‰, a decrease to a Permian minimum of
about +10‰ and an increase towards the modern
ocean value of +20‰ (Claypool 1980; Strauss 1997).
Superimposed over this long-term secular variation is
a pronounced short-term variability exhibiting about
the same isotopic range as documented for the entire
Phanerozoic. The observed variations are assumed to
reflect changes in the sulfur redox cycle within the
ocean. Since bacterial sulfate reduction and the subsequent formation of sedimentary pyrite is associated
with a substantial fractionation, any changes in the
isotopic mass balance are generally interpreted as
changes in the burial of reduced sulfur. Furthermore,
both the rate of sulfur input and its isotopic composition may vary as a function of time depending on the
weathering rates of different sedimentary rocks, or on
the intensity of volcanic activity. For example, increased erosion of black shales containing 34 S-depleted sulfides may lower the δ 34 S value of oceanic
sulfate. On the other hand, the increased reduction of
sulfate by bacteria and removal of 34 S-depleted
sulfide from the ocean may increase its δ 34 S value.
However, the fact that the deposition of evaporites in
the Phanerozoic history occurred in rather short
episodes raises the question about the validity of the
sulfate isotope curve with respect to representativeness and homogeneity. The specific causes for the
variation of δ 34 S of marine sulfate at a particular
geologic time interval are still largely speculative (for
further discussion see Strauss 1997).
A special case in the modern ocean sulfur isotope distribution is the isotopic composition of hydrogen sulfide in anoxic basins. Fry et al. (1991)
compared depth profiles of the Black Sea and the