345
δ 18 O PDB = 0.97002 · δ 18 O SMOW - 29.98
(10.7)
A δ 18 O w correction of –0.27‰ is therefore necessary to compare δ 18 O values measured in CO 2
produced by the reaction of calcite with H 3 PO 4
with those measured in CO 2 equilibrated with
water.
Most of the paleotemperature equations appear to be similar to the one stated above (Eq.
10.6), but temperature reconstructions can differ
as much as 2°C when ambient temperature varies
between 5°C and 25°C (Fig. 10.2; Bemis et al.
1998; for a recent review see Mulitza et al. 2003).
The reason is that, in addition to in-situ temperature and water isotopic composition, the shell
δ 18 O may be affected by the photosynthetic
activity of algal symbionts and by the carbonate
ion concentration in seawater. Wefer and Berger
(1991) summarized the importance of such effects, previously called 'vital effects', on a broad
spectrum of organisms (Fig. 10.3). For oxygen
isotopes, most organisms are shown to precipitate CaCO 3 close to equilibrium with the water in
which they live. However, some organisms, such
as planktonic foraminifera and hermatypic corals,
exhibit significant differences from isotopic equilibrium (see Waelbroek et al. 2005 for a review).
Laboratory experiments with live planktonic
foraminifera demonstrated that an increase in the
symbiont photosynthetic activity results in a
decrease in shell δ 18 O values (Spero 1992; Spero
and Lea 1993). The mechanisms driving the
effects of symbiont photosynthesis on shell δ 18 O
are not well understood, but appear to be linked
to the carbonate ion concentration. During
photosynthetic activity, CO 2 uptake by the
symbionts increases the pH in the microenvironment around the shell (Jørgensen et al.
1985; Rink 1998; Zeebe 1999). Consequently, more
alkaline conditions correspond to locally
elevated [CO 3
2- ]. The 18 O depletion of shells due
to higher symbiont photosynthetic activity is
consistent with the effect of higher ambient
[CO 3
2- ]. Spero et al. (1997) obtained a δ 18 O/[CO 3
2- ]
slope of –0.002‰ µmol -1 kg -1 from experiments
with symbiotic (O. universa ) and non-symbiotic
(G. bulloides) foraminifer species. This previously undocumented carbonate isotope effect
may help to solve inconsistencies in temperature
reconstructions by applying oxygen isotope
paleothermometry relative to other marine or terrestrial temperature proxies (see discussions in
Spero et al. 1997; Zeebe 1999).
Diagenesis
The isotopic composition of a carbonatic shell will
remain unchanged until the shell material dissolves
and recrystallizes during diagenesis. Diagenetic
modification, however, could begin immediately
after deposition or even in the water column due
to corrosive deep ocean or pore waters. Such waters are generally enriched in CO 2 due to the respiration of organic matter mediated by specific bacteria. In shells of the planktonic foraminifera P.
obliquiloculata sampled in a depth profile in the
western equatorial Pacific Wu and Berger (1989)
showed that below the depth of the modern
lysocline the δ 18 O of this species increased with
water depth due to increasing calcite dissolution.
Close to the modern depth of calcium carbonate
compensation, the observed deviation reached a
maximum value of +0.9‰. This effect of differential dissolution, i.e. the preferential removal of the
light isotope 16 O, has recently helped to unravel
the so-called cool tropics paradox. Using
exceptionally well preserved planktonic foraminifer shells extracted from clay-rich late Cretaceous
to Eocene sediments, Pearson et al. (2001) could
show that the previously reported cooler tropical
sea surface temperatures are the result of up to 3
‰ increased δ
18
Ο values caused by a subtle
diagenetic recrystallization on a micrometer scale,
without obliterating surface ornaments or internal
layering features. Pearson et al. (2001) emphasized that this kind of diagenesis is hard to
detect compared to more obvious modes of alteration like the precipitation of euhedral inorganic
calcite resulting in substantial infillings or overgrowth.
The conversion of sediment into limestone
within deep-sea sediments results from pressure
and temperature rises which both increase with
burial depth. Within carbonate sediments, diagenesis generally transforms the less stable
aragonite and Mg-calcite into a low-Mg calcitic
cement by means of a dissolution-reprecipitation
process. Theoretically, the oxygen isotope composition of carbonates should not change significantly with burial, since the 18 O in pore waters
originates from seawater. However, in many cores
recovered during DSDP/ODP drilling, deep-sea
carbonates and often pore waters as well, exhibit
18 O depletions by several permil (e.g. Lawrence
1989). Mass balance calculations by Matsumoto
(1992) indicated that the 18 O shift in pore water towards lower values in sediments of the Japan Sea
10.3
Geochemical Influences on 18 O/ 16 O Ratios
δ 18 O PDB = 0.97002 · δ 18 O SMOW - 29.98
(10.7)
A δ 18 O w correction of –0.27‰ is therefore necessary to compare δ 18 O values measured in CO 2
produced by the reaction of calcite with H 3 PO 4
with those measured in CO 2 equilibrated with
water.
Most of the paleotemperature equations appear to be similar to the one stated above (Eq.
10.6), but temperature reconstructions can differ
as much as 2°C when ambient temperature varies
between 5°C and 25°C (Fig. 10.2; Bemis et al.
1998; for a recent review see Mulitza et al. 2003).
The reason is that, in addition to in-situ temperature and water isotopic composition, the shell
δ 18 O may be affected by the photosynthetic
activity of algal symbionts and by the carbonate
ion concentration in seawater. Wefer and Berger
(1991) summarized the importance of such effects, previously called 'vital effects', on a broad
spectrum of organisms (Fig. 10.3). For oxygen
isotopes, most organisms are shown to precipitate CaCO 3 close to equilibrium with the water in
which they live. However, some organisms, such
as planktonic foraminifera and hermatypic corals,
exhibit significant differences from isotopic equilibrium (see Waelbroek et al. 2005 for a review).
Laboratory experiments with live planktonic
foraminifera demonstrated that an increase in the
symbiont photosynthetic activity results in a
decrease in shell δ 18 O values (Spero 1992; Spero
and Lea 1993). The mechanisms driving the
effects of symbiont photosynthesis on shell δ 18 O
are not well understood, but appear to be linked
to the carbonate ion concentration. During
photosynthetic activity, CO 2 uptake by the
symbionts increases the pH in the microenvironment around the shell (Jørgensen et al.
1985; Rink 1998; Zeebe 1999). Consequently, more
alkaline conditions correspond to locally
elevated [CO 3
2- ]. The 18 O depletion of shells due
to higher symbiont photosynthetic activity is
consistent with the effect of higher ambient
[CO 3
2- ]. Spero et al. (1997) obtained a δ 18 O/[CO 3
2- ]
slope of –0.002‰ µmol -1 kg -1 from experiments
with symbiotic (O. universa ) and non-symbiotic
(G. bulloides) foraminifer species. This previously undocumented carbonate isotope effect
may help to solve inconsistencies in temperature
reconstructions by applying oxygen isotope
paleothermometry relative to other marine or terrestrial temperature proxies (see discussions in
Spero et al. 1997; Zeebe 1999).
Diagenesis
The isotopic composition of a carbonatic shell will
remain unchanged until the shell material dissolves
and recrystallizes during diagenesis. Diagenetic
modification, however, could begin immediately
after deposition or even in the water column due
to corrosive deep ocean or pore waters. Such waters are generally enriched in CO 2 due to the respiration of organic matter mediated by specific bacteria. In shells of the planktonic foraminifera P.
obliquiloculata sampled in a depth profile in the
western equatorial Pacific Wu and Berger (1989)
showed that below the depth of the modern
lysocline the δ 18 O of this species increased with
water depth due to increasing calcite dissolution.
Close to the modern depth of calcium carbonate
compensation, the observed deviation reached a
maximum value of +0.9‰. This effect of differential dissolution, i.e. the preferential removal of the
light isotope 16 O, has recently helped to unravel
the so-called cool tropics paradox. Using
exceptionally well preserved planktonic foraminifer shells extracted from clay-rich late Cretaceous
to Eocene sediments, Pearson et al. (2001) could
show that the previously reported cooler tropical
sea surface temperatures are the result of up to 3
‰ increased δ
18
Ο values caused by a subtle
diagenetic recrystallization on a micrometer scale,
without obliterating surface ornaments or internal
layering features. Pearson et al. (2001) emphasized that this kind of diagenesis is hard to
detect compared to more obvious modes of alteration like the precipitation of euhedral inorganic
calcite resulting in substantial infillings or overgrowth.
The conversion of sediment into limestone
within deep-sea sediments results from pressure
and temperature rises which both increase with
burial depth. Within carbonate sediments, diagenesis generally transforms the less stable
aragonite and Mg-calcite into a low-Mg calcitic
cement by means of a dissolution-reprecipitation
process. Theoretically, the oxygen isotope composition of carbonates should not change significantly with burial, since the 18 O in pore waters
originates from seawater. However, in many cores
recovered during DSDP/ODP drilling, deep-sea
carbonates and often pore waters as well, exhibit
18 O depletions by several permil (e.g. Lawrence
1989). Mass balance calculations by Matsumoto
(1992) indicated that the 18 O shift in pore water towards lower values in sediments of the Japan Sea
10.3
Geochemical Influences on 18 O/ 16 O Ratios