10 Influence of Geochemical Processes on Stable Isotope Distribution in Marine Sediments
356
removal of particulate nitrogen from the system, the
change of δ 15 N within this accumulating pool is
described by Equation 10.22, the integral of Equations 10.20 and 10.21. However, this latter case
almost never occurs (Altabet and Francois 1994).
A topic of much concern regarding the application of bulk δ 15 N in paleoceanographic investigations is the effect of terrestrially derived organic
matter on the marine δ 15 N signal. Since terrigenous
material is mostly substantially depleted in 15 N
relative to marine-derived matter, the nitrogen
isotopic composition of ocean sediments have been
used to trace the origin of sedimentary nitrogen
(Sweeny and Kaplan 1980; Wada et al. 1987).
However, the underlying assumption regarding the
mixing of a marine and terrestrial δ 15 N end members
is not always conclusive because of the wide range
of δ 15 N values produced in terrestrial systems
(Sweeney et al. 1978). Furthermore, marine values
may be subject to changes produced by processes
in the water column (as discussed above).
Diagenesis
Reconstructing past variations of nutrient conditions in surface waters requires the preservation of
the nitrogen isotope surface-water signal in the
sedimentary organic matter. Diagenesis, however,
may significantly alter the nitrogen isotopic composition of particles even before they are buried at the
seafloor. Altabet and McCarthy (1985) suggested
that bacterial remineralization of sinking particles
resulted in a 15 N enrichment in the residual organic
matter. Francois et al. (1992) reported a 15 N
enrichment of 5‰ to 7‰ in core top values of
Southern Ocean sediments compared to measurements of organic matter in the photic zone. However,
Altabet and Francois (1994) demonstrated that there
was only little offset between sinking particles at
150 m water depth in the equatorial Pacific and the
respective core tops. Sigman et al. (1999) suggested
that the difference in preservation of the isotopic
signal existent in the two regions may be explained
with the high opal content in the Southern Ocean
sediments, in which organic matrices with high δ 15 N
may be captured. In any case, the findings of these
authors suggest that where differences between the
surface-generated signal and sediments exist, the
offset appears to be relatively constant within geographic regions.
Further decomposition of organic matter occurs
at the sediment/water interface and in the upper
sections of the sediment. However, observations
from equatorial Pacific sediments indicate that the
loss of organic carbon as well as of nitrogen in the
uppermost few millimeters is not accompanied by a
corresponding shift in δ 15 N over this depth interval
(Altabet and Francois 1994). Freudenthal et al.
(2001) found that early diagenesis in sediments of
the eastern tropical Atlantic led to variability in
δ 15 N of even less than 1‰. With further diagenesis
of organic matter, nitrogen may occur as ammonium
incorporated in the lattice of clay minerals, where it
replaces potassium. But even this complex
substitution is not associated with an isotopic
effect, since this nitrogen exhibits an isotopic
composition very similar to the organic matter from
which it is derived (Williams 1995). Therefore,
beside the possible fractionation mechanisms,
which may alter the nitrogen isotopic composition
of organic matter in the water column, there seems
to be no significant diagenetic alteration of the
isotope signal within the sediments. The nitrogen
isotopic composition of sediments is primarily determined by the source organic matter.
10.6 Geochemical Influences
on 34 S / 32 S Ratios
10.6.1 δ
δ δ
δ δ 34 S of Seawater and Pore Waters
Principles of Fractionation
In the marine environment, sulfur occurs most
commonly in its oxidized form as dissolved sulfate
in seawater or as precipitated sulfate in evaporites
and in its reduced form as sedimentary pyrite. The
ratio of 34 S/ 32 S is a sensitive indicator for the
Fig. 10.10 Main reservoirs of the sedimentary sulfur
cycle (adopted from Strauss 1997).
356
removal of particulate nitrogen from the system, the
change of δ 15 N within this accumulating pool is
described by Equation 10.22, the integral of Equations 10.20 and 10.21. However, this latter case
almost never occurs (Altabet and Francois 1994).
A topic of much concern regarding the application of bulk δ 15 N in paleoceanographic investigations is the effect of terrestrially derived organic
matter on the marine δ 15 N signal. Since terrigenous
material is mostly substantially depleted in 15 N
relative to marine-derived matter, the nitrogen
isotopic composition of ocean sediments have been
used to trace the origin of sedimentary nitrogen
(Sweeny and Kaplan 1980; Wada et al. 1987).
However, the underlying assumption regarding the
mixing of a marine and terrestrial δ 15 N end members
is not always conclusive because of the wide range
of δ 15 N values produced in terrestrial systems
(Sweeney et al. 1978). Furthermore, marine values
may be subject to changes produced by processes
in the water column (as discussed above).
Diagenesis
Reconstructing past variations of nutrient conditions in surface waters requires the preservation of
the nitrogen isotope surface-water signal in the
sedimentary organic matter. Diagenesis, however,
may significantly alter the nitrogen isotopic composition of particles even before they are buried at the
seafloor. Altabet and McCarthy (1985) suggested
that bacterial remineralization of sinking particles
resulted in a 15 N enrichment in the residual organic
matter. Francois et al. (1992) reported a 15 N
enrichment of 5‰ to 7‰ in core top values of
Southern Ocean sediments compared to measurements of organic matter in the photic zone. However,
Altabet and Francois (1994) demonstrated that there
was only little offset between sinking particles at
150 m water depth in the equatorial Pacific and the
respective core tops. Sigman et al. (1999) suggested
that the difference in preservation of the isotopic
signal existent in the two regions may be explained
with the high opal content in the Southern Ocean
sediments, in which organic matrices with high δ 15 N
may be captured. In any case, the findings of these
authors suggest that where differences between the
surface-generated signal and sediments exist, the
offset appears to be relatively constant within geographic regions.
Further decomposition of organic matter occurs
at the sediment/water interface and in the upper
sections of the sediment. However, observations
from equatorial Pacific sediments indicate that the
loss of organic carbon as well as of nitrogen in the
uppermost few millimeters is not accompanied by a
corresponding shift in δ 15 N over this depth interval
(Altabet and Francois 1994). Freudenthal et al.
(2001) found that early diagenesis in sediments of
the eastern tropical Atlantic led to variability in
δ 15 N of even less than 1‰. With further diagenesis
of organic matter, nitrogen may occur as ammonium
incorporated in the lattice of clay minerals, where it
replaces potassium. But even this complex
substitution is not associated with an isotopic
effect, since this nitrogen exhibits an isotopic
composition very similar to the organic matter from
which it is derived (Williams 1995). Therefore,
beside the possible fractionation mechanisms,
which may alter the nitrogen isotopic composition
of organic matter in the water column, there seems
to be no significant diagenetic alteration of the
isotope signal within the sediments. The nitrogen
isotopic composition of sediments is primarily determined by the source organic matter.
10.6 Geochemical Influences
on 34 S / 32 S Ratios
10.6.1 δ
δ δ
δ δ 34 S of Seawater and Pore Waters
Principles of Fractionation
In the marine environment, sulfur occurs most
commonly in its oxidized form as dissolved sulfate
in seawater or as precipitated sulfate in evaporites
and in its reduced form as sedimentary pyrite. The
ratio of 34 S/ 32 S is a sensitive indicator for the
Fig. 10.10 Main reservoirs of the sedimentary sulfur
cycle (adopted from Strauss 1997).