measured in oceanic regions close to upwelling of intermediate and deep waters in the North Indian and Pacific
Oceans. DR is the local departure from the global mean
value of R. The latter is calculated from the modeled global
sea surface
14 C age, using a global box-diffusion carbon
model, which accounts for the
14 C production and carbon
cycle (Reimer et al. 2013).
In the past, the oceanic circulation changed as the reservoir ages did. Changes of the differences between the marine
and atmospheric
14 C ages may be estimated by dating the
charcoals and foraminifera from widespread volcanic ash
layers in both marine sediments and on land (Bard et al.
1994; Siani et al. 2001; Austin et al. 2011; Thornalley et al.
2011; Siani et al. 2013; Sikes and Guilderson 2016), as well
as by dating paired charcoals and mollusk shells found in the
same deposits (Bondevik et al. 2006; Ascough et al. 2009).
These studies demonstrated that sea surface R changed
during the first step of the deglaciation from the modern
value of 400 to 2000 years in the North Atlantic (Bard et al.
1994; Bondevik et al. 2006; Austin et al. 2011; Thornalley
et al. 2011) and to 800 years in the Mediterranean Sea (Siani
et al. 2013). In the South Pacific, the subtropical R augmented from *300 to *700 years (Sikes and Guilderson
2016), while those of the sub Antarctic surface waters
increased from *800 to 1400 years and to 3200 years
(Siani et al. 2013; Sikes and Guilderson 2016).
Past increases in the sea ice extent at high latitudes during
the glacial periods prevents the atmosphere-ocean
14 CO 2
exchanges contributing to the increase of the sea surface R.
In the IntCal13 calibration record, the modern value of R
was subtracted to the
14 C ages of the marine samples to be
compared to the atmospheric
14 C ages. To account for
changes in the oceanic circulation, the modern value of R
was augmented by 200 years from *14,000 cal BP and
then considered as constant down to 50,000 cal BP (Reimer
et al. 2013).
Continental Environments: The Hard Water
and Dead Carbon Effects
Modern lake vegetation and calcium carbonates in lakes in
calcareous regions exhibit older
14 C ages than those of the
atmosphere. The dissolved inorganic carbon
14 CO 2 in lakes,
used during photosynthetic processes and during the precipitation of calcium carbonates originates from the dissolution of
14 C-free carbonates of geological age (dead carbon)
and from the mineralization of old organic matter enclosed
in lake sediments. Impact of the later on the apparent
14 C
ages of lacustrine plants is much higher in lake surrounded
by peats and in artificial lakes implemented by soils flooding. The resulting aging of
14 C ages is called ‘hard water’
effect. In addition, water stratification in lakes or the presence of an ice cover in high altitude and polar lakes tend to
prevent the atmospheric
14 CO 2 input to the lake waters, that
tends to increase the
14 C age in lake waters relatively to that
of the atmosphere.
The speleothems are composed of calcium carbonates,
formed by the dissolution of
14 C-free geological carbonates
by slightly acidified waters by CO 2 from the atmosphere and
from the degradation of the soil organic matter. In Fig. 4.7
the procedure to estimate the fraction of dead carbon
(DCF) in a Bahamas speleothem, located in the western
North Atlantic, is shown (Beck et al. 2001). To an U-Th age
corresponds an atmospheric
14 C age in the calibration
record. The DCF is calculated by subtracting the
14 C age of
the speleothem to that of the contemporaneous atmosphere.
In this example, the
14 C ages of the speleothem are older
than those of the atmosphere at an average of about
1450 years between 11,000 cal. BP and 15,000 cal. BP,
which corresponds to a DCF of 16%. This aging is not
constant as a function of time and varies between 1000 years
and 2000 years. These variations appear to be closely correlated to the climatic fluctuations recorded either in the
Greenland ice-core GISP2 or in the marine varved sediments
of the Cariaco Basin. About 30% of the variability of the
DCF in this speleothem may be explained by such fluctuations by the way of changes of the local temperature and
rainfall patterns.
The
14
C Exchanges in the Carbon Reservoirs
Radiocarbon is commonly used to test numerical simulations
of the oceanic circulation in Ocean General Circulation
Models (O-GCM) (Toggweiler et al. 1989; Key et al. 2004).
The systematic measurements of the
14 C content of the
dissolved inorganic carbon (DIC) in the different world
ocean basins started with the international oceanographic
campaigns GEOSECS (1972-1978), followed by WOCE
(World Ocean Experiment, 1990–2002), and others
(Broecker et al. 1995; Key et al. 1996). In addition to
14 C
analyses, the physicochemical properties of the worldwide
basin water masses were measured along depth profiles.
Both the natural and anthropogenic components of the
14 C
concentration in the oceans offer the opportunity to validate
the general ocean circulation simulated by numerical models
(Toggweiler et al. 1989; Key et al. 2004). The natural
component tests the circulation of deep waters, while the
anthropogenic component, resulting from the thermonuclear
tests in the 1960s, allows the analysis of physical processes
with time constants of a few decades, such as the formation
of deep and intermediate waters and the ventilation of the
thermocline (transition zone between the cold intermediate
and deep waters and the warm surface waters). Because of
radioactive decay, the
14 C content of an ocean water mass
decreases during the oceanic transport, once the water mass
is isolated from exchanges with the atmosphere.
4 Carbon-14
61
Oceans. DR is the local departure from the global mean
value of R. The latter is calculated from the modeled global
sea surface
14 C age, using a global box-diffusion carbon
model, which accounts for the
14 C production and carbon
cycle (Reimer et al. 2013).
In the past, the oceanic circulation changed as the reservoir ages did. Changes of the differences between the marine
and atmospheric
14 C ages may be estimated by dating the
charcoals and foraminifera from widespread volcanic ash
layers in both marine sediments and on land (Bard et al.
1994; Siani et al. 2001; Austin et al. 2011; Thornalley et al.
2011; Siani et al. 2013; Sikes and Guilderson 2016), as well
as by dating paired charcoals and mollusk shells found in the
same deposits (Bondevik et al. 2006; Ascough et al. 2009).
These studies demonstrated that sea surface R changed
during the first step of the deglaciation from the modern
value of 400 to 2000 years in the North Atlantic (Bard et al.
1994; Bondevik et al. 2006; Austin et al. 2011; Thornalley
et al. 2011) and to 800 years in the Mediterranean Sea (Siani
et al. 2013). In the South Pacific, the subtropical R augmented from *300 to *700 years (Sikes and Guilderson
2016), while those of the sub Antarctic surface waters
increased from *800 to 1400 years and to 3200 years
(Siani et al. 2013; Sikes and Guilderson 2016).
Past increases in the sea ice extent at high latitudes during
the glacial periods prevents the atmosphere-ocean
14 CO 2
exchanges contributing to the increase of the sea surface R.
In the IntCal13 calibration record, the modern value of R
was subtracted to the
14 C ages of the marine samples to be
compared to the atmospheric
14 C ages. To account for
changes in the oceanic circulation, the modern value of R
was augmented by 200 years from *14,000 cal BP and
then considered as constant down to 50,000 cal BP (Reimer
et al. 2013).
Continental Environments: The Hard Water
and Dead Carbon Effects
Modern lake vegetation and calcium carbonates in lakes in
calcareous regions exhibit older
14 C ages than those of the
atmosphere. The dissolved inorganic carbon
14 CO 2 in lakes,
used during photosynthetic processes and during the precipitation of calcium carbonates originates from the dissolution of
14 C-free carbonates of geological age (dead carbon)
and from the mineralization of old organic matter enclosed
in lake sediments. Impact of the later on the apparent
14 C
ages of lacustrine plants is much higher in lake surrounded
by peats and in artificial lakes implemented by soils flooding. The resulting aging of
14 C ages is called ‘hard water’
effect. In addition, water stratification in lakes or the presence of an ice cover in high altitude and polar lakes tend to
prevent the atmospheric
14 CO 2 input to the lake waters, that
tends to increase the
14 C age in lake waters relatively to that
of the atmosphere.
The speleothems are composed of calcium carbonates,
formed by the dissolution of
14 C-free geological carbonates
by slightly acidified waters by CO 2 from the atmosphere and
from the degradation of the soil organic matter. In Fig. 4.7
the procedure to estimate the fraction of dead carbon
(DCF) in a Bahamas speleothem, located in the western
North Atlantic, is shown (Beck et al. 2001). To an U-Th age
corresponds an atmospheric
14 C age in the calibration
record. The DCF is calculated by subtracting the
14 C age of
the speleothem to that of the contemporaneous atmosphere.
In this example, the
14 C ages of the speleothem are older
than those of the atmosphere at an average of about
1450 years between 11,000 cal. BP and 15,000 cal. BP,
which corresponds to a DCF of 16%. This aging is not
constant as a function of time and varies between 1000 years
and 2000 years. These variations appear to be closely correlated to the climatic fluctuations recorded either in the
Greenland ice-core GISP2 or in the marine varved sediments
of the Cariaco Basin. About 30% of the variability of the
DCF in this speleothem may be explained by such fluctuations by the way of changes of the local temperature and
rainfall patterns.
The
14
C Exchanges in the Carbon Reservoirs
Radiocarbon is commonly used to test numerical simulations
of the oceanic circulation in Ocean General Circulation
Models (O-GCM) (Toggweiler et al. 1989; Key et al. 2004).
The systematic measurements of the
14 C content of the
dissolved inorganic carbon (DIC) in the different world
ocean basins started with the international oceanographic
campaigns GEOSECS (1972-1978), followed by WOCE
(World Ocean Experiment, 1990–2002), and others
(Broecker et al. 1995; Key et al. 1996). In addition to
14 C
analyses, the physicochemical properties of the worldwide
basin water masses were measured along depth profiles.
Both the natural and anthropogenic components of the
14 C
concentration in the oceans offer the opportunity to validate
the general ocean circulation simulated by numerical models
(Toggweiler et al. 1989; Key et al. 2004). The natural
component tests the circulation of deep waters, while the
anthropogenic component, resulting from the thermonuclear
tests in the 1960s, allows the analysis of physical processes
with time constants of a few decades, such as the formation
of deep and intermediate waters and the ventilation of the
thermocline (transition zone between the cold intermediate
and deep waters and the warm surface waters). Because of
radioactive decay, the
14 C content of an ocean water mass
decreases during the oceanic transport, once the water mass
is isolated from exchanges with the atmosphere.
4 Carbon-14
61