347
ocean and the Atlantic deep water. The observed
deviations occur because net biological production in the Southern Ocean is low relative to the
upwelling fluxes of nutrients and dissolved carbon, allowing low 13 C/ 12 C ratios in the dissolved
carbon to outcrop and exchange with the atmosphere. As a consequence, a dynamic balance is
achieved where CO 2 evading from Antarctic waters to the atmosphere has a lower δ 13 C than the
invading CO 2 , while the isotope ratio of the evading CO 2 from the warm ocean is greater than that
of the return flux of CO 2 .
Modern Range of Values and Historical Variability
The modern δ 13 C ΣCO2 values of seawater are close
to 0‰ (PDB) and vary only within a small range.
From the GEOSECS δ 13 C ΣCO2 sections for today’s
world oceans, compiled by Kroopnick (1985), a
range of +2.5‰ in the mid-latitude surface Atlantic to +0.7‰ in the northern surface Pacific has
been determined. Deep water mass δ 13 C ΣCO2
ranges from +1.2‰ in the core of North Atlantic
Deep Water and +0.4‰ in the Circumpolar Deep
Water to –1.0‰ in the northern Pacific Deep Water. As stated above, δ 13 C ΣCO2 of a deep water
mass behaves like a conservative tracer for phosphate, since the water mass left its sea surface
source area. On the other hand, paired δ 13 C ΣCO2
and PO 4 measurements on two sections in the
Southern Ocean clearly support the thermodynamic imprint during Weddell Sea Bottom Water
formation by a significant deviation from the
Redfield δ 13 C ΣCO2 /PO 4 relationship (Mackensen et
al. 1996; Mackensen 2001). A recent matter of
concern is the so-called oceanic Suess effect,
which describes the uptake of
13
C-depleted
anthropogenic CO 2 by the oceans and the
consequent decrease in the oceanic δ
13
C.
Comparisons between modern sediment trap and
core top foraminiferal δ
13
C reveal a modern
depletion in δ
13
C of the upper water column of up
to 0.65‰ in the Subantarctic Zone and up to
0.9‰ in the Arctic realm (e.g., King and Howard
2004). This effect has to be considered before
using surface sediment values for carbon cycle
proxy calibrations.
The δ 13 C ΣCO2 varied considerably in geologic
history. Generally, three explanations are given for
changes of δ 13 C ΣCO2 distribution in the ocean: (1)
changes in the surface-ocean productivity which
cause variable fractionation between surface and
deep water carbon isotopic composition, (2)
changes in the gas exchange rates between ocean
and atmosphere due to changes in surface
temperatures and ocean circulation, and (3)
changes in the marine carbon budget by variations in the reservoirs of the atmosphere, the
ocean, or the lithosphere. Of course, these different processes act on different time scales. Variations on a scale of 10 1 to 10 5 years might be attributed to changes in the effectivity of the biologic
pump (Berger and Vincent 1986), to rapid changes
between land and ocean reservoirs (Shackleton
1977; Broecker 1982), or to changes in ocean circulation responding to climate variability (e.g.
Sarnthein et al. 1994; Bickert and Mackensen 2003).
On scales of 10 5 years and longer, changes in
the influx and in the burial of sedimentary inorganic and organic carbon exert a primary control
on the δ 13 C of the ocean and the atmosphere
(Holser 1997). Main sources for carbon comprise
the erosional flux of sedimentary carbon (as C org
and carbonate) and the degassing of volcanic
CO 2 ; main sinks are the burial of organic matter
and the deposition of carbonate. Changes in the
oceanic δ 13 C are assumed to be the result of variations in the ratio of inorganic to organic carbon
contributed to sediments (e.g. Derry and FranceLanord 1996). An increase in the burial of organic
carbon would preferentially remove 12 C from
seawater, so that the ocean reservoir would become isotopically heavier, and vice versa.
10.4.2 δ
δ δ
δ δ 13 C in Marine Organic Matter
Principles of Fractionation
The carbon isotopic composition of marine organic matter produced in the photic zone depends
on the isotope ratio of the total dissolved inorganic carbon and the degree to which this inorganic pool is utilized (Degens et al. 1968). Although more than 90% of the inorganic carbon
pool is in the form of HCO 3
- , phytoplankton utilizes mostly carbon of the very small reservoir
(1%) of dissolved CO 2(aq) . While the equilibrium
exchange fractionation is only small between atmospheric CO 2 and dissolved CO 2(aq) (about 1‰),
and moderate between CO 2(aq) and HCO 3
- (with a
range of 10‰ to 8‰ between 5°C and 25°C; Mook
et al. 1974), a large kinetic fractionation accompanies the biological carbon fixation during photosynthesis. The main isotope-discriminating steps
are (1) the uptake and intracellular diffusion of
CO 2 , and (2) the enzymatic carbon fixation (Park
10.4
Geochemical Influences on 13 C/ 12 C Ratios
ocean and the Atlantic deep water. The observed
deviations occur because net biological production in the Southern Ocean is low relative to the
upwelling fluxes of nutrients and dissolved carbon, allowing low 13 C/ 12 C ratios in the dissolved
carbon to outcrop and exchange with the atmosphere. As a consequence, a dynamic balance is
achieved where CO 2 evading from Antarctic waters to the atmosphere has a lower δ 13 C than the
invading CO 2 , while the isotope ratio of the evading CO 2 from the warm ocean is greater than that
of the return flux of CO 2 .
Modern Range of Values and Historical Variability
The modern δ 13 C ΣCO2 values of seawater are close
to 0‰ (PDB) and vary only within a small range.
From the GEOSECS δ 13 C ΣCO2 sections for today’s
world oceans, compiled by Kroopnick (1985), a
range of +2.5‰ in the mid-latitude surface Atlantic to +0.7‰ in the northern surface Pacific has
been determined. Deep water mass δ 13 C ΣCO2
ranges from +1.2‰ in the core of North Atlantic
Deep Water and +0.4‰ in the Circumpolar Deep
Water to –1.0‰ in the northern Pacific Deep Water. As stated above, δ 13 C ΣCO2 of a deep water
mass behaves like a conservative tracer for phosphate, since the water mass left its sea surface
source area. On the other hand, paired δ 13 C ΣCO2
and PO 4 measurements on two sections in the
Southern Ocean clearly support the thermodynamic imprint during Weddell Sea Bottom Water
formation by a significant deviation from the
Redfield δ 13 C ΣCO2 /PO 4 relationship (Mackensen et
al. 1996; Mackensen 2001). A recent matter of
concern is the so-called oceanic Suess effect,
which describes the uptake of
13
C-depleted
anthropogenic CO 2 by the oceans and the
consequent decrease in the oceanic δ
13
C.
Comparisons between modern sediment trap and
core top foraminiferal δ
13
C reveal a modern
depletion in δ
13
C of the upper water column of up
to 0.65‰ in the Subantarctic Zone and up to
0.9‰ in the Arctic realm (e.g., King and Howard
2004). This effect has to be considered before
using surface sediment values for carbon cycle
proxy calibrations.
The δ 13 C ΣCO2 varied considerably in geologic
history. Generally, three explanations are given for
changes of δ 13 C ΣCO2 distribution in the ocean: (1)
changes in the surface-ocean productivity which
cause variable fractionation between surface and
deep water carbon isotopic composition, (2)
changes in the gas exchange rates between ocean
and atmosphere due to changes in surface
temperatures and ocean circulation, and (3)
changes in the marine carbon budget by variations in the reservoirs of the atmosphere, the
ocean, or the lithosphere. Of course, these different processes act on different time scales. Variations on a scale of 10 1 to 10 5 years might be attributed to changes in the effectivity of the biologic
pump (Berger and Vincent 1986), to rapid changes
between land and ocean reservoirs (Shackleton
1977; Broecker 1982), or to changes in ocean circulation responding to climate variability (e.g.
Sarnthein et al. 1994; Bickert and Mackensen 2003).
On scales of 10 5 years and longer, changes in
the influx and in the burial of sedimentary inorganic and organic carbon exert a primary control
on the δ 13 C of the ocean and the atmosphere
(Holser 1997). Main sources for carbon comprise
the erosional flux of sedimentary carbon (as C org
and carbonate) and the degassing of volcanic
CO 2 ; main sinks are the burial of organic matter
and the deposition of carbonate. Changes in the
oceanic δ 13 C are assumed to be the result of variations in the ratio of inorganic to organic carbon
contributed to sediments (e.g. Derry and FranceLanord 1996). An increase in the burial of organic
carbon would preferentially remove 12 C from
seawater, so that the ocean reservoir would become isotopically heavier, and vice versa.
10.4.2 δ
δ δ
δ δ 13 C in Marine Organic Matter
Principles of Fractionation
The carbon isotopic composition of marine organic matter produced in the photic zone depends
on the isotope ratio of the total dissolved inorganic carbon and the degree to which this inorganic pool is utilized (Degens et al. 1968). Although more than 90% of the inorganic carbon
pool is in the form of HCO 3
- , phytoplankton utilizes mostly carbon of the very small reservoir
(1%) of dissolved CO 2(aq) . While the equilibrium
exchange fractionation is only small between atmospheric CO 2 and dissolved CO 2(aq) (about 1‰),
and moderate between CO 2(aq) and HCO 3
- (with a
range of 10‰ to 8‰ between 5°C and 25°C; Mook
et al. 1974), a large kinetic fractionation accompanies the biological carbon fixation during photosynthesis. The main isotope-discriminating steps
are (1) the uptake and intracellular diffusion of
CO 2 , and (2) the enzymatic carbon fixation (Park
10.4
Geochemical Influences on 13 C/ 12 C Ratios