Tracing of seawater masses, surface water
productivity, and ecology
The d
13
C DIC in water is widely used as a tracer for seawater masses (e.g., Ravelo and Hillarie-Marcel, 2007).
Although changes in surface water productivity and
mixing influence d
13 C DIC , other processes, including
water mass formation and ocean circulation, determine
the d
13
C DIC composition and therefore can be used as
tracers for water masses. Deep water captures the d
13
C
signature from its surface water sources. North Atlantic
Deep Water (NADW) has therefore relatively high d
13
C
values (~1.11 %), whereas Antarctic Bottom Water
(ABW) is relatively low (~0.4 %), reflecting the major
differences in surface water d
13
C DIC in both high-latitude
source regions (Kroopnick, 1985). Variations in deep
water d
13
C can therefore be used to trace the history of
deep water currents. This approach is also referred to as
“aging” of deep water (Rohling and Cook, 1999), where
the relative age of deep water provides a measure for
ocean circulation.
The d
13
C difference between shallow and deepdwelling planktic foraminifera (Dd
13
C) provides a wellestablished proxy for thermocline productivity and the
depth of the shallow-water mixed layer (e.g., Wefer
et al., 1999). Increasing Dd
13
C values reflect more stratified surface water masses, due to increasing contrasts
between shallow and deep waters and vice versa.
A comparable approach is used to reconstruct productivity
via changes in benthic foraminifera Dd
13
C of epifaunal
and infaunal species (Zahn et al., 1986). Here, the d
13
C
differences are mainly controlled by pore water geochemical gradients and the intensity of organic matter degradation, the latter being a function of organic matter supply
(flux) to the sea floor (McCorkle et al., 1990). Large
Dd
13
C values represent less organic matter flux and thus
less primary productivity and vice versa. Benthic foraminifera Dd
13
C gradients can, however, be biased by the variability in new production of organic matter (fresh
phytodetritus), which influences the stable isotope signal.
This effect, constrained by variable primary productivity
and organic matter fluxes at seasonal time scales, is
referred to as the “phytodetritus” or “Mackensen effect”
(Mackensen et al., 1993). Furthermore, the carbon and
oxygen isotopic composition of calcite tests of planktic
foraminifera is used for reconstructing depth habitats and
thus the ecology of extant and extinct species (e.g., Pearson et al., 1993).
Chemostratigraphy and carbon isotope
excursions (CIE)
The history of the global carbon cycle and climate through
time is preserved in the carbon isotope record of marine
and terrestrial sediments (Figure 2). The application of
organic and inorganic carbon isotope ratios as chemical
fingerprints for stratigraphic correlations was pioneered
in the late 1970s with the studies of Berger et al. (1978),
Weissert et al. (1979), and Scholle and Arthur (1980),
more recently summarized in Weissert et al. (2008),
Saltzman and Thomas (2012), and Wendler (2013). The
most robust Phanerozoic archives of marine carbon isotope signatures are from individual pristine preserved fossils and bulk carbonate from marine pelagic sediments
(e.g., Joachimski and Buggisch, 1993; Veizer et al.,
1999; Saltzman and Thomas, 2012; and references
therein). Carbon isotope fluctuations of several per mil
are also reported from hemipelagic sediments, shallowwater carbonates, and terrestrial sediments. Due to multiple local influences, however, they may differ in their
absolute isotope values and ranges from their
corresponding pelagic sediment records, complicating
their use in stratigraphic correlation across wider areas.
The long-term trend in Earth’s climate is interrupted by
series of transient events, identified by distinct positive or
negative carbon isotope excursions, CIEs (Figure 2).
These CIEs, combined with longer trends in the global
carbon isotope record, are used in chemostratigraphy to
identify chemical events (e.g., Weissert et al., 1998;
Zachos et al., 2001), which can be correlated within and
across ocean basins and the terrestrial environment (e.g.,
Gröcke et al., 1999; Herrle et al., 2004). Well-documented
examples of CIEs representing major perturbations of the
global carbon cycle are reported from, e.g., the PaleogeneMesozoic greenhouse (Paleocene-Eocene thermal maximum PETM, Cretaceous and Early Jurassic oceanic
anoxic events, OAEs, Jenkyns, 2003) and the Paleozoic
(Saltzman and Thomas, 2012; Figure 2). Many CIEs coincide with widespread environmental perturbations including severe and short-term global warming (hyperthermal
events), ocean acidification and shoaling of the carbonate
compensation depth (CCD), enhanced marine organic carbon burial linked to widespread ocean anoxia, major crisis
or extinction of biota on land and in the ocean, and sudden
shifts in Earth’s hydrological cycle and climate. Understanding how the Earth system responded to and recovered from these past extreme events remains a major
focal point of Earth sciences, adding to the discussion on
anthropogenically induced global warming.
The magnitude and rapidity of CIEs in the Phanerozoic,
combined with evidence for massive environmental
change, have stimulated intense interdisciplinary research
about the underlying trigger and feedback mechanisms.
Positive CIEs are reported for multimillion to orbital timescales and are, among other possible mechanisms
(Figure 2), commonly associated with periods of globally
enhanced marine carbon burial, leading to the concept of
oceanic anoxic events (Schlanger and Jenkyns, 1976;
Scholle and Arthur, 1980) and global cooling (e.g.,
Kuypers et al., 1999). Recent studies, combining geochemical with biotic data and biogeochemical modeling,
have shown that global cooling can cause perturbations
to marine ecosystems and biogeochemical cycles at scales
comparable to those associated with global warming
(McAnena et al., 2013). Sharp negative CIEs, either isolated within the chemostratigraphic record or in combination with positive CIEs, have been recorded throughout
CARBON ISOTOPES
75
productivity, and ecology
The d
13
C DIC in water is widely used as a tracer for seawater masses (e.g., Ravelo and Hillarie-Marcel, 2007).
Although changes in surface water productivity and
mixing influence d
13 C DIC , other processes, including
water mass formation and ocean circulation, determine
the d
13
C DIC composition and therefore can be used as
tracers for water masses. Deep water captures the d
13
C
signature from its surface water sources. North Atlantic
Deep Water (NADW) has therefore relatively high d
13
C
values (~1.11 %), whereas Antarctic Bottom Water
(ABW) is relatively low (~0.4 %), reflecting the major
differences in surface water d
13
C DIC in both high-latitude
source regions (Kroopnick, 1985). Variations in deep
water d
13
C can therefore be used to trace the history of
deep water currents. This approach is also referred to as
“aging” of deep water (Rohling and Cook, 1999), where
the relative age of deep water provides a measure for
ocean circulation.
The d
13
C difference between shallow and deepdwelling planktic foraminifera (Dd
13
C) provides a wellestablished proxy for thermocline productivity and the
depth of the shallow-water mixed layer (e.g., Wefer
et al., 1999). Increasing Dd
13
C values reflect more stratified surface water masses, due to increasing contrasts
between shallow and deep waters and vice versa.
A comparable approach is used to reconstruct productivity
via changes in benthic foraminifera Dd
13
C of epifaunal
and infaunal species (Zahn et al., 1986). Here, the d
13
C
differences are mainly controlled by pore water geochemical gradients and the intensity of organic matter degradation, the latter being a function of organic matter supply
(flux) to the sea floor (McCorkle et al., 1990). Large
Dd
13
C values represent less organic matter flux and thus
less primary productivity and vice versa. Benthic foraminifera Dd
13
C gradients can, however, be biased by the variability in new production of organic matter (fresh
phytodetritus), which influences the stable isotope signal.
This effect, constrained by variable primary productivity
and organic matter fluxes at seasonal time scales, is
referred to as the “phytodetritus” or “Mackensen effect”
(Mackensen et al., 1993). Furthermore, the carbon and
oxygen isotopic composition of calcite tests of planktic
foraminifera is used for reconstructing depth habitats and
thus the ecology of extant and extinct species (e.g., Pearson et al., 1993).
Chemostratigraphy and carbon isotope
excursions (CIE)
The history of the global carbon cycle and climate through
time is preserved in the carbon isotope record of marine
and terrestrial sediments (Figure 2). The application of
organic and inorganic carbon isotope ratios as chemical
fingerprints for stratigraphic correlations was pioneered
in the late 1970s with the studies of Berger et al. (1978),
Weissert et al. (1979), and Scholle and Arthur (1980),
more recently summarized in Weissert et al. (2008),
Saltzman and Thomas (2012), and Wendler (2013). The
most robust Phanerozoic archives of marine carbon isotope signatures are from individual pristine preserved fossils and bulk carbonate from marine pelagic sediments
(e.g., Joachimski and Buggisch, 1993; Veizer et al.,
1999; Saltzman and Thomas, 2012; and references
therein). Carbon isotope fluctuations of several per mil
are also reported from hemipelagic sediments, shallowwater carbonates, and terrestrial sediments. Due to multiple local influences, however, they may differ in their
absolute isotope values and ranges from their
corresponding pelagic sediment records, complicating
their use in stratigraphic correlation across wider areas.
The long-term trend in Earth’s climate is interrupted by
series of transient events, identified by distinct positive or
negative carbon isotope excursions, CIEs (Figure 2).
These CIEs, combined with longer trends in the global
carbon isotope record, are used in chemostratigraphy to
identify chemical events (e.g., Weissert et al., 1998;
Zachos et al., 2001), which can be correlated within and
across ocean basins and the terrestrial environment (e.g.,
Gröcke et al., 1999; Herrle et al., 2004). Well-documented
examples of CIEs representing major perturbations of the
global carbon cycle are reported from, e.g., the PaleogeneMesozoic greenhouse (Paleocene-Eocene thermal maximum PETM, Cretaceous and Early Jurassic oceanic
anoxic events, OAEs, Jenkyns, 2003) and the Paleozoic
(Saltzman and Thomas, 2012; Figure 2). Many CIEs coincide with widespread environmental perturbations including severe and short-term global warming (hyperthermal
events), ocean acidification and shoaling of the carbonate
compensation depth (CCD), enhanced marine organic carbon burial linked to widespread ocean anoxia, major crisis
or extinction of biota on land and in the ocean, and sudden
shifts in Earth’s hydrological cycle and climate. Understanding how the Earth system responded to and recovered from these past extreme events remains a major
focal point of Earth sciences, adding to the discussion on
anthropogenically induced global warming.
The magnitude and rapidity of CIEs in the Phanerozoic,
combined with evidence for massive environmental
change, have stimulated intense interdisciplinary research
about the underlying trigger and feedback mechanisms.
Positive CIEs are reported for multimillion to orbital timescales and are, among other possible mechanisms
(Figure 2), commonly associated with periods of globally
enhanced marine carbon burial, leading to the concept of
oceanic anoxic events (Schlanger and Jenkyns, 1976;
Scholle and Arthur, 1980) and global cooling (e.g.,
Kuypers et al., 1999). Recent studies, combining geochemical with biotic data and biogeochemical modeling,
have shown that global cooling can cause perturbations
to marine ecosystems and biogeochemical cycles at scales
comparable to those associated with global warming
(McAnena et al., 2013). Sharp negative CIEs, either isolated within the chemostratigraphic record or in combination with positive CIEs, have been recorded throughout
CARBON ISOTOPES
75