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GEOPHYSICAL RESEARCH LETTERS, VOL. 34, L21705, doi:10.1029/2007GL031304, 2007
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Younger Dryas: A data to model comparison to constrain the strength
of the overturning circulation
Katrin J. Meissner1
Received 12 July 2007; revised 24 September 2007; accepted 9 October 2007; published 8 November 2007.
[1] The University of Victoria Earth System Climate
Model (UVic ESCM) is used to compare simulated time
series of radiocarbon during the Younger Dryas (YD) with
paleoceanographic records. I find that only a complete shutdown and recovery of the Atlantic Meridional Overturning
Circulation (AMOC) can simulate both the rise in
atmospheric CO2 concentrations seen in ice core records
and the peak and subsequent decrease in atmospheric D14C
comparable to the peak recorded in the varved sediments of
the Cariaco Basin. Simulated radiocarbon profiles in the
western North Atlantic match well with data from deep-sea
corals at the beginning of the YD, whereas planktonic/
benthic foraminifera records match best with a transient
state during the rapid recovery of the AMOC. The steepness
of the increase in atmospheric D14C at !12.9 ka cal could
not be simulated with oceanic circulation changes only
because the response time of the climate system is too slow.
Citation: Meissner, K. J. (2007), Younger Dryas: A data to
model comparison to constrain the strength of the overturning
circulation, Geophys. Res. Lett., 34, L21705, doi:10.1029/
2007GL031304.
1. Introduction
[2] The Younger Dryas (YD, 12.9 – 11.5 cal ka) is one of
the three major abrupt cold events that interrupted the
warming of the Northern Hemisphere during the last deglaciation. It can be recognized in a variety of tracers in
terrestrial [e.g., Mathewes et al., 1993], marine [e.g.,
Broecker et al., 1989] and ice records [e.g., Alley et al.,
1993] across the Northern Hemisphere. Johnson and
McClure [1976] first postulated that an increased freshwater
flux to the St. Lawrence River associated with the diversion
of continental runoff from the Mississippi River triggered
the YD cold event by causing an increase in North Atlantic
sea ice. Rooth [1982] subsequently proposed that this
Mississippi-to-St. Lawrence routing caused the YD through
its effect on the thermohaline circulation and attendant
poleward ocean heat transport.
[3] The YD is a relatively recent event. Radiocarbon
(D14C) is therefore a valuable proxy that can help us to
understand the trigger and mechanism of this abrupt cooling. 14C is produced at the top of the atmosphere where
neutrons interact with nitrogen atoms. The production rate is
highly dependent on solar activity and the strength of the
terrestrial magnetic field [Muscheler et al., 2004]. After
production, 14C oxidizes rapidly to 14CO2 and is incorpo1
School of Earth and Ocean Sciences, University of Victoria, Victoria,
Canada.
Copyright 2007 by the American Geophysical Union.
0094-8276/07/2007GL031304$05.00
rated in the global carbon cycle. Since radiocarbon decays
with a half-life of 5730 years [Godwin, 1962], carbon
reservoirs which have been isolated from the atmosphere
for a long time (such as the deep ocean) contain considerably lower amounts of 14C than ‘‘newer’’ carbon reservoirs.
Changes in the oceanic circulation and ventilation will
therefore alter the atmosphere-ocean 14C flux. The atmospheric D14C is influenced both by production at the top of
the atmosphere and the exchange with ‘‘old’’ carbon
reservoirs.
[4] Atmospheric D14C records are compiled from paleo
proxies in lake sediments [Goslar et al., 1995, 1999], corals
[Fairbanks et al., 2005; Chiu et al., 2007] and varved
sediments of the Cariaco Basin [Hughen et al., 2000,
2004a, 2004b]. The tree-ring chronology [Stuiver et al.,
1998; Friedrich et al., 2004] extends back to 12.40 cal ka
and is therefore too young to capture the signature of the
whole YD event. Radiocarbon is also used to reconstruct
shifts in water masses and evaluate possible changes in the
oceanic meridional heat transport. D14C in deep-sea corals
[Robinson et al., 2005; Eltgroth et al., 2006] and the age
difference between benthic and planktonic foraminifera
[Keigwin, 2004] give an insight about the change in
ventilation and water masses at a certain location.
[5] In the last ten years, numerous studies have been
published in an attempt to understand what caused the
variability in the atmospheric radiocarbon record during
the YD [e.g. Beck et al., 2001]. The researchers divide into
two schools: those who emphasize on changes in the
production rate [e.g., Muscheler et al., 2000; Marchal et
al., 2001; Goslar et al., 2000] and those who highlight
evidence for a change in the oceanic circulation [e.g.,
Goslar et al., 1995; Hughen et al., 1998; Clark et al.,
2002; Adkins and Pasquero, 2004; McManus et al., 2004].
Radiocarbon was also integrated into a variety of physical
or carbon cycle models: Stocker and Wright [1996] used a
zonally averaged physical ocean model and found an
increase in atmospheric D14C of 35 per mill when the
Atlantic Meridional Overturning Circulation (AMOC) shuts
down. With a newer version of the same model, Marchal et
al. [2001] could not explain the atmospheric D14C peak
with either production or circulation changes. On the other
hand, several carbon cycle box models have been
employed: Goslar et al. [1999] concluded using the
PANDORA box model that variations in geomagnetic field
are too weak to explain the D14C maximum seen in the
records. The simulations of Köhler et al. [2005] using the
BICYCLE box model showed that the modelled pCO2 and
d13C concentrations were most consistent with data when
only a small change in ocean circulation was imposed; they
concluded that the D14C record must be caused by changes
in the production rate.
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Figure 1. (a) Freshwater flux added to the St. Lawrence
River Delta (blue) and strength of the overturning motion
for run A (red) and B (green) in Sv. Note that the blue
shaded line is only valid for simulation B. The colored stars
indicate the times of the D14C profiles shown in Figure 2.
(b) Time series of global carbon budgets (minus mean value
for each reservoir) in PgC for run B only, atmosphere
(black), ocean (cyan), terrestrial carbon (orange).
[6] Here I present the results of a state-of-the-art three
dimensional ocean model, coupled to an atmosphere, sea
ice, land surface and vegetation model when forced with a
freshwater event. The model incorporates a fully coupled
carbon cycle and prognostic D14C. The results are compared to the atmospheric D14C record [Hughen et al.,
2004a], radiocarbon in benthic foraminifera [Keigwin,
2004] and deep-sea corals [Eltgroth et al., 2006], in order
to constrain the strength of the AMOC during the YD event.
I concentrate on the effect circulation changes have on the
atmospheric D14C and keep therefore the 14C production
rate constant during my simulations.
2. Model and Experimental Design
[7] The University of Victoria (UVic) Earth System
Climate Model (ESCM, version 2.8) consists of an ocean
general circulation model (Modular Ocean Model version 2
[Pacanowski, 1995]) coupled to a vertically integrated two
dimensional energy-moisture balance model of the atmo-
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sphere, a dynamic-thermodynamic sea ice model based on
the works by Semtner [1976], Hibler [1979] and Hunke and
Dukowitz [1997], a land surface scheme (simplified version
of ‘‘MOSES’’ [Cox et al., 1999]) and a dynamic global
vegetation model (‘‘TRIFFID’’ [Meissner et al., 2003b;
Cox, 2001]). The model including the atmospheric, ocean
and sea ice components is described by Weaver et al.
[2001]. It is driven by seasonal variations in solar insolation
at the top of the atmosphere and seasonally-varying wind
stress and wind fields [Kalnay et al., 1996]. The UVic
ESCM also includes a fully coupled carbon cycle taking
into account the terrestrial carbon fluxes and reservoirs
[Meissner et al., 2003b; Matthews et al., 2005] as well as
the inorganic [Ewen et al., 2004] and organic [Schmittner et
al., 2007] carbon cycle in the ocean. Radiocarbon has been
added as a prognostic tracer and is treated in carbon
reservoirs and carbon fluxes like 12C with the only difference being that it decays with time. Present day control
simulations compare well to GEOSECS data [Meissner et
al., 2003a].
[8] I integrate the model for over 12000 years into
equilibrium under boundary conditions corresponding to
12.9 ka cal (orbital parameters of 12.9 ka cal [Berger,
1978], atmospheric CO2 of 242 ppm and elevated topography [Peltier, 1994]). During this equilibrium run, I hold the
atmospheric D14C fixed at 200 per mill. At the end of the
equilibrium run, I diagnose the 14C production rate at the
top of the atmosphere needed to keep the atmospheric
14 12
C/ C ratio constant. This procedure yields a production
rate of 1.61 atoms/(cm2s).
[9] The imposed meltwater event follows the reconstruction of Carlson et al. [2007] and consists of four different
regimes in the discharge from the St. Lawrence River (see
Figure 1): 0.06 Sv is added from 12.9 to 12.5 k cal; 0.12 Sv
from 12.5 to 12.3 k cal; 0 Sv from 12.3 to 11.9 k cal and
0.07 Sv from 11.9 to 11.5 k cal. During the simulations of the
meltwater event, the 14C production rate is held constant and
atmospheric 14C and CO2 are therefore calculated prognostically. Two simulations are carried out, one in which the
freshwater scenario is set equal to the discharge from the St.
Lawrence River (run A, red line in Figure 1a) and one in
which the freshwater is added to the simulated discharge
from St. Lawrence River (run B, green line in Figure 1a).
The difference in freshwater forcing between run A and run
B depends on the amount of precipitation over land (and
therefore on the climate state); its average equals 0.03 Sv.
3. Results
[10] Figure 1a shows the time series of the freshwater
scenario (blue, solid line) as well as the maximum strength
of the meridional overturning for runs A and B. Starting
with a maximum strength of 18 Sv, the red simulation
(run A) shows a weakening of 36% (11.5 Sv) followed by a
short recovery and another weakening to 12.5 Sv. At the end
of the event, run A recovers to its initial strength. Run B on
the other hand shows a complete shut down of the AMOC
and does not recover on its own after the freshwater event.
As there is paleoceanographic data evidence that the thermohaline circulation recovered after the YD, I apply an
unphysical salt flux of "0.12 Sv to the St. Lawrence River
delta during simulation B in order to simulate CO2 and
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Figure 2. Simulated D14C profiles at 30!N, 70!W compared to Keigwin’s [2004] profile (black stars, connected by black
line, error bars in gray) and Eltgroth et al.’s [2006] deep-sea coral profile (pink circles connected by pink line).
(a) Snapshots at 12.9, 12.8, 12.6, 12.2 and 11.7 k cal. (b) Snapshots at 11.4, 11.3, 11.2 and 11.1 k cal (see also Figure 1).
D14C during a recovery (Figure 1a, blue dashed line). Run
B reacts instantly with a recovery to the initial strength of
the meridional overturning first and a further increase to a
maximum of 35 Sv.
[11] The global carbon budgets for run B are shown in
Figure 1b. Carbon is transferred from the atmosphere (black
line) and land (orange line) into the ocean (cyan line) during
the time the AMOC is shut down. With the recovery of the
AMOC deep water enriched in dissolved organic carbon
(DIC) is brought to the surface where a net flux of oceanic
inorganic carbon into the atmosphere takes place [Meissner
et al., 2007]. This carbon accumulates in the atmosphere
and results in an increase of the atmospheric CO2 concentration above initial conditions. With warmer conditions in
the Northern Hemisphere due to an enhanced oceanic heat
transport and higher atmospheric CO2 concentrations, vegetation and soil carbon also increases significantly.
[12] D14C profiles in the western North Atlantic have
been reconstructed using benthic/planktonic foraminifera
[Keigwin, 2004] and deep-sea corals [Eltgroth et al.,
2006] and are shown in Figure 2. The most striking
characteristic of Keigwin’s [2004] profile is the strong
contrast in ventilation at !2.3 km depth between well
ventilated Glacial North Atlantic Intermediate Water and
poorly ventilated glacial deep water. Also shown in Figure 2
are snapshots of D14C profiles from my B simulation at
30!N, 70!W. The time of each snapshot is indicated in
Figure 1 using colored stars (the colors correspond to the
colors of the profiles in Figure 2). None of the simulated
profiles match the high D14C values of Keigwin [2004]
above 2.3 km depth. The shut down of the AMOC results in
a gradual decrease of D14C in the deep ocean and a slight
increase of D14C at the surface due to a longer residence
time and therefore more effective gas exchange (Figure 2a).
With the recovery of the AMOC, intermediate water masses
get more ventilated and the positive D14C anomaly propagates downward (Figure 2b). This is the only time where
the model captures a strong D14C gradient at intermediate
depth as indicated by foraminifera data. The last profile (in
blue) shows that the gradient in D14C around 2000 to 3000
meter depths completely disappeared once the AMOC has
recovered and the ocean had time to re-equilibrate. This
profile is also similar to the first profile (in red) before the
event. The two data points of the deep-sea corals (pink
circles [Eltgroth et al., 2006]) match best with my first two
profiles at the beginning of the YD. This is consistent with
the corals’ age (!12.8 ka). Very recent data from the
California Margin [Marchitto et al., 2007] show a depletion
of radiocarbon in intermediate waters during the YD. My
simulations do not simulate a similar drop at the same
location (23.5!N, 111.6!W). As our ocean model operates
on a coarse grid (3.6! # 1.8!) with simplified topography, it
is not able to simulate local shifts in circulation so close to
shore.
[13] Figure 3a shows the atmospheric D14C record
[Hughen et al., 2004a] compared to the simulated atmospheric D14C during run A (red) and B (green, solid line). A
weakening of the AMOC (run A) is unable to trigger
significant changes in the atmospheric D14C. Run B
(green), however, shows a peak at 244 per mill and a
subsequent decline which is similar to the data. The most
important difference between model and data is due to the
response time of the climate system in my model. Although
the AMOC collapses within 300 years, atmospheric D14C
does not react instantly and builds up slowly over the whole
period the ventilation is reduced. Once the AMOC recovers,
simulated atmospheric D14C declines at the same rate as
shown in the paleo data (Figure 3a, green dashed line).
The model response in D14C lags the reconstruction by
900 years, indicating that the YD freshwater event might
have been shorter than the scenario used for the present
simulations.
[14] Finally, Figure 3b compares the CO2 concentrations
from EPICA Dome C ice cores [Monnin et al., 2001] with
the simulated atmospheric CO2 concentrations (run A in
red; run B in green, solid lines). As for D14C the model’s
response lags the reconstruction by several hundred years
pointing to an earlier recovery of the AMOC than possible
with the meltwater scenario chosen for the present simulations (the green dashed line shows the response of run B
shifted back by 900 years). Although the resumption of the
AMOC provokes a rise in atmospheric CO2 which is similar
to the initial rise in CO2 seen in data, the simulated CO2
levels out around 260 ppm.
4. Conclusions
[15] By incorporating radiocarbon as a prognostic tracer
in the UVic ESCM and presuming that the production rate
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rable to ice core data [Monnin et al., 2001]. Radiocarbon
profiles in the western North Atlantic match well with data
from deep-sea corals at the beginning of the YD [Eltgroth et
al., 2006], whereas planktonic/benthic foraminifera records
[Keigwin, 2004] match best with a transient response during
the rapid recovery of the AMOC from a complete shutdown state.
[16] Acknowledgments. The author would like to thank Michael Eby
and Ed Wiebe for their technical support, as well as two anonymous
reviewers for their helpful and constructive comments. Katrin J. Meissner is
grateful for research grant support under the University Faculty Award
programm (NSERC).
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Figure 3. (a) Time series of simulated atmospheric D14C
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green dashed line corresponds to the green simulation
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AMOC in our simulation. (b) Time series of simulated
atmospheric CO2 concentrations (runs A and B in red and
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Dome C (black dots [Monnin et al., 2001]). The green
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atmospheric CO2 concentrations with the resumption of
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K. J. Meissner, School of Earth and Ocean Sciences, University of
Victoria, PO Box 3055, Stn CSC, Victoria, BC, Canada, V8W 3P6.
([email protected])
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