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Author's personal copy
Quaternary Science Reviews 30 (2011) 1155e1172
Contents lists available at ScienceDirect
Quaternary Science Reviews
journal homepage: www.elsevier.com/locate/quascirev
Deconstructing the Last Glacial termination: the role of millennial
and orbital-scale forcings
L. Menviel a, b, *, A. Timmermann c, O. Elison Timm c, A. Mouchet d
a
Climate and Environmental Physics, Physics Institute, University of Bern, Sidlerstrasse 5, CH-3012 Bern, Switzerland
Oeschger Centre for Climate Change Research, University of Bern, CH-3012 Bern, Switzerland
c
IPRC, SOEST, University of Hawai’i, 2525 Correa Road, Honolulu, HI 96822, USA
d
Département AGO, Université de Liège, Liège, Belgium
b
a r t i c l e i n f o
a b s t r a c t
Article history:
Received 29 September 2010
Received in revised form
15 February 2011
Accepted 16 February 2011
Available online 11 March 2011
Using an Earth system model of intermediate complexity forced by continuously varying boundary
conditions and a hypothetical profile of freshwater forcing, the model simulates Heinrich event 1 (H1),
the Bølling warm period, the Older Dryas, the Antarctic Cold Reversal (ACR) and the Younger Dryas in
close agreement with paleo-proxy data from different regions worldwide. The ACR can be simulated as
the bipolar seesaw response to the AMOC recovery during the termination of H1. However, this study
also demonstrates that the amplitude of the ACR can be further amplified by a rapid deglacial retreat of
the Antarctic Ice sheets. We suggest that melting from both, the Laurentide and the Antarctic Ice sheets
contributed to the sea level rise associated with Meltwater Pulse 1-A (MWP-1A). It is hypothesized that
the northern hemispheric source of MWP-1A caused the Older Dryas cooling in the Northern Hemisphere, whereas the Southern Hemispheric source contributed to the ACR. The study also documents that
for the majority of paleo-climate proxies considered here, the relative timing can be qualitatively
reproduced by the transient modeling experiments. The climate model solution presented here may
provide a means to further constrain dating uncertainties of some of paleo-climate proxies during the
Last Glacial Termination.
Ó 2011 Elsevier Ltd. All rights reserved.
Keywords:
Last Glacial termination
Heinrich event 1
Bølling warming
Older Dryas
Antarctic Cold Reversal
Younger Dryas
MWP-1A
Earth system model of intermediate
complexity
1. Introduction
Elucidating the driving mechanisms of glacial cycles requires
both an accurate knowledge of the timing and magnitude of
external forcings and an understanding of internal climate feedbacks that determine the overall climate response to these forcings.
In addition, paleo-climate reconstructions that can be used to
independently test the hypothesized mechanisms are essential for
ascertaining our comprehension of these changes.
Paleo-modeling efforts have played an important role in quantifying the climate response to particular glacial-scale forcings.
Using coupled atmosphere-ocean models, these efforts have mostly
focused on particular time-slices, such as the Last Glacial Maximum
(LGM, 21 ka B.P.) (Bush and Philander, 1998; Liu et al., 2000, 2002;
Hewitt et al., 2001; Kitoh et al., 2001; Kim et al., 2003; Shin et al.,
2003; Timmermann et al., 2004; Braconnot et al., 2007) or the
* Corresponding author. Climate and Environmental Physics, Physics Institute,
University of Bern, Sidlerstrasse 5, CH-3012 Bern, Switzerland.
E-mail address: [email protected] (L. Menviel).
0277-3791/$ e see front matter Ó 2011 Elsevier Ltd. All rights reserved.
doi:10.1016/j.quascirev.2011.02.005
Mid-Holocene Optimum (9e6 ka B.P.) (Joussaume et al., 1999; OttoBliesner, 1999; Liu et al., 2003). Results from these timeslice
experiments have been compared thoroughly with existing paleoreconstructions, such as CLIMAP-Project-Members (1981), GLAMAP (Pflaumann et al., 2003) and other proxies (see e.g. Shin et al.
(2003) for a careful LGM model-data comparison). This modeling
approach has helped to explore the spatial characteristics of past
glacial climate change and understand the relative contributions of
different external forcing factors (Rind, 1987; Timmermann et al.,
2004; Justino et al., 2005).
However, in order to understand the physical mechanisms
responsible for abrupt climate change and to identify the existence
of leads, lags and phase-relationships in the climate system, it is
mandatory to simulate the temporal evolution of the system in
response to the time-varying forcing factors. Recently the transient
modeling framework has been applied to climate models to
understand the Late Quaternary climate evolution and its relation
to slowly-varying boundary conditions (Jackson and Broccoli, 2003;
Felis et al., 2004; Lorenz and Lohmann, 2004; Charbit et al., 2005;
Tuenter et al., 2005; Renssen et al., 2005; Lorenz et al., 2006;
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L. Menviel et al. / Quaternary Science Reviews 30 (2011) 1155e1172
Lunt et al., 2006; Marsh et al., 2006; Timmermann et al., 2007,
2009a; Timm and Timmermann, 2007; Liu et al., 2009;
Timmermann and Menviel, 2009). We recently presented a transient model solution for the Last Glacial Termination using the
Earth system model of intermediate complexity LOVECLIM that was
driven by changes in greenhouse gases, icesheet topography and
albedo and orbitally-induced changes in solar radiation (Timm and
Timmermann, 2007). Millennial-scale freshwater forcing was
however not considered. More recently, Liu et al. (2009) performed
a transient simulation for the period 22 to 15 ka B.P. using a state of
the art CGCM (NCAR CCSM3) that also captured Heinrich event 1
(H1) and the transition to the Bølling warm period (Timmermann
and Menviel, 2009).
The meltdown of the Laurentide and Eurasian ice sheets started
around 20 ka B.P. An acceleration occurred during H1 (17.8e15.2 ka
B.P.), as massive icebergs broke off from the disintegrating ice
sheets. The resulting discharge of freshwater into the North Atlantic
(Vidal et al., 1997) reduced surface density and eventually halted
the formation of North Atlantic Deep Water for about 2500 years, as
suggested by the 231Pa/230Th and ventilation age data from North
Atlantic sediment cores (McManus et al., 2004; Gherardi et al.,
2005; Thornalley et al., 2011) (Fig. 1a). At the same time, temperatures over Greenland remained low (Alley, 2000) while temperatures over Antarctica started to rise (Jouzel et al., 2007) and
atmospheric CO2 concentration increased from 190 to 220 ppmv
(Monnin et al., 2001). The end of H1 is marked by an abrupt
warming over Greenland and the North Atlantic starting around
14.7 ka B.P. Liu et al. (2009) suggested that this Bølling warming
was due to the rapid resumption of the AMOC following H1,
combined with the glacial-interglacial increase in atmospheric CO2.
Fig. 1. (a) Time series of the maximum of the meridional stream function (Sv) in the North Atlantic for experiment DGNS (green) compared to 231Pa/230Th data (black) from core
OCE326-GGC5 (McManus et al., 2004) (filled circles) and from core SU81-18 (Gherardi et al., 2005) (empty circles) as well as ventilation age data (gray) from cores RAPID 10-1P, 154P and 17-5P (Thornalley et al., 2011). The time series of (McManus et al., 2004) was shifted by 200 years backward to better fit the GISP2 data (Alley, 2000). H1 stands for Heinrich
event 1, BWP for Bolling Warm Period, OD for Older Dryas and YD for Younger Dryas. (b) Time series of the maximum overturning strength in the North Pacific for experiment DGNS
(green) compared to the ventilation ages (yrs) as recorded in marine sediment cores from the Western North Pacific (black). The ventilation ages curve is the smoothed spline
interpolation of averaged benthiceplanktic foraminifera ages and projection ages in the Western North Pacific (Okazaki et al., 2010). (c) Time series of dissolved oxygen content (ml/
L) averaged over the California basin (135 W-114 W, 25 N-40 N) for experiment DGNS (green) compared to a composite of two d15N record (black) from the California Basin
(marine sediment cores JPC-56 (Pride et al., 1999) and ODP 893A (Emmer and Thunell, 2000) and compared to the Mo/Al data from core ODP 893A (gray circles) (Ivanochko and
Pedersen, 2004). (d) Freshwater input (Sv) in the North Atlantic (black) and in the Southern Ocean (green) as applied in experiment DGNS. Time series are 21 year running means for
the model outputs. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)
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L. Menviel et al. / Quaternary Science Reviews 30 (2011) 1155e1172
Subsequent to the Bølling warming, between 14.7 and 13 ka B.P.,
the temperature over Greenland showed some pronounced fluc
tuations of about 5 C within 400e500 years. Two major cold events
were identified during this period in the Norwegian Sea (Lehman
and Keigwin, 1992) and in Greenland (Stuiver et al., 1995): the
Older Dryas (w14 ka B.P.) and the intra-Allerød cold period
(13.5e13 ka B.P.) (Fig. 2a).
During the same time, a 2 C surface cooling, known as the
Antarctic Cold Reversal (ACR) is reconstructed for Antarctica (Petit
et al., 1999; Jouzel et al., 2007). Concurrent to the ACR, sea level rose
by about 20 m in about 500 years (Fairbanks, 1989; Stanford et al.,
2006). The origin of this rapid sea level rise, referred to as Meltwater Pulse 1A (MWP-1A), has been the subject of an intense
debate. While Licht (2004) and Peltier (2005) propose a Northern
Hemispheric origin, Kanfoush et al. (2000); Clark et al. (2002);
Bassett et al. (2005) and Carlson (2009) advocate the idea of
a Southern Hemispheric origin of MWP-1A. Evidence for
1157
a meltwater pulse originating from Antarctica at around 14.6 ka B.P.
was found in several marine sediment cores of the Southern Ocean.
Ice-Rafted Debris (IRD) layers were observed in sediment cores
from the Southeast Atlantic (Kanfoush et al., 2000), the Southwest
Pacific (Carter et al., 2002) and from around the Antarctic Peninsula
(Noumi et al., 2007). Freshwater discharge in the Southern Ocean is
also inferred from depleted oxygen values in planktic foraminifera
as well as increased occurrence of marine diatoms (Chaetoceros
spp) in a sediment core from the Atlantic sector of the Southern
Ocean (Bianchi and Gersonde, 2004). Recent studies by Clark et al.
(2002) and Bassett et al. (2005) further suggest that in order to
reconcile coral-based sea level reconstructions during MWP-1A
with predicted glacial isostatic adjustment, one has to invoke
a significant melting of Antarctica. Bassett et al. (2005) estimated
that the Antarctic icesheet could have contributed by up to about
15 m sea level rise during MWP-1A. The main potential source
regions for this pulse in Antarctica were identified as the Weddell
Fig. 2. (a) Time series of spring season (MAM) air temperature anomalies ( C) averaged over Greenland (47 W-0 , 60 N-80 N) for experiment DGNS (green) compared to the
temperature reconstruction from Greenland ice core GISP2 (black) (Alley, 2000). IACP stands for Intra-Allerød Cold Period. (b) Summer (JJA) air temperature anomalies ( C) averaged
over the Netherlands (0 E10 E, 45 N-55 N) for experiment DGNS (green) compared to chironomid-inferred July air temperature anomalies ( C) from Hijkermeer, the Netherlands
(Heiri et al., 2007). (c) SST anomalies ( C) averaged over the Iberian margin (11 W-8 W, 37 N-40 N) for experiment DGNS (green) compared to a reconstructed temperature index
(black) obtained by averaging different alkenone-based SST reconstructions from cores SU81-18 (Bard et al., 2000), MD95-2042 (Bard, 2002), ODP-161-977A (Martrat et al., 2004)
and MD01-2444 (Martrat et al., 2007). The time axis of the alkenone-based SST reconstructions was shifted forward by 200 years to better fit the model results and the other
northern hemispheric paleo-proxies. (d) Winter season (DJF) SST anomalies ( C) averaged over the Cariaco basin (55 W-52 W, 10 N-14 N) for experiment DGNS (green) compared
to Mg/Ca SST reconstruction ( C, black) from core PL07-39PC (Lea et al., 2003). As marine productivity is the highest during the winter season, this season was chosen to best
represent temperature variations. Anomalies are with respect to pre-industrial time. Time series are 21 year running means for the model outputs. (For interpretation of the
references to colour in this figure legend, the reader is referred to the web version of this article.)
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L. Menviel et al. / Quaternary Science Reviews 30 (2011) 1155e1172
Sea (w9 m) and the Ross Sea (w5 m). Finally, recent modeling
studies (Philippon et al., 2006; Pollard and DeConto, 2009) suggest
that the accelerated melting of the Antarctic icesheet started
around 15 ka, further supporting the notion of significant deglacial
freshwater forcing of the Southern Ocean during the ACR. The ACR
could be explained in terms of the bipolar seesaw response to the
AMOC recovery during the Bølling warming (Stocker, 1998), but
there is also the possibility that a near-synchronous rapid melting
of the Antarctic icesheet may have contributed to the Southern
Hemispheric cooling (Menviel et al., 2010a).
The Younger Dryas (YD) transition to colder Northern Hemispheric climates occurred around 12.9 ka B.P. (Rasmussen et al.,
2006). During this period temperatures over Greenland, the
eastern North Atlantic and the Caribbean Sea dropped by 8 C, 2 C
and 4 C, respectively (Fig. 2). It has been hypothesized that the YD
was caused by a weakening of the AMOC (McManus et al., 2004)
due to an outburst of the late glacial Lake Agassiz into the Arctic
Ocean (Tarasov and Peltier, 2005; Murton et al., 2010). Its effect on
the Southern Hemisphere still remains somewhat elusive.
Our goal is to simulate the Last Glacial Termination using an
Earth system model of intermediate complexity, LOVECLIM, and to
address the following questions:
What forcings are needed to reproduce the dominant climate
and biogeochemical features of the Last Glacial Termination?
What are the global impacts of H1, MWP-1A and the YD?
Are climate data consistent with the notion of a Southern
Hemispheric origin of MWP-1A?
Is it justified to view the North Atlantic as the main driver of
millennial-scale variability during the Last Glacial
Termination?
Is the relative timing of events in paleo-proxy data from
different regions consistent with the relative timing in the
model?
A suite of transient climate modeling experiments is conducted
for the period 18e11 ka B.P. using time-varying radiative forcing
due to orbitally-induced solar radiation changes and greenhouse
gases. The effects of waning glacial ice sheets on topography and
albedo are captured by updating the atmospheric boundary
conditions continuously with estimates of the icesheet evolution
and albedo. Millennial-scale variability is produced by varying the
strength of the deep water formation in the North Atlantic and in
the Southern Ocean through the input of idealized freshwater
pulses.
The paper is organized as follows: in section 2 we provide
a description of the experimental setup and the methods employed
to analyze the paleo-proxy data. Section 3 is devoted to an extensive comparison between model simulation and paleo-climate
proxies covering the period 18e11 ka B.P. The paper concludes with
a general discussion and a summary of the main results.
2. Model and experimental setup
2.1. The Earth system model LOVECLIM
The model used in this study is the model of intermediate
complexity LOVECLIM (Driesschaert et al., 2007; Goosse et al.,
2007)1. Our experiments were conducted with version 1.1. of this
model. The atmospheric component of the coupled model LOVECLIM is ECBilt (Opsteegh et al., 1998), a spectral T21, three-level
model, based on quasi-geostrophic equations extended by
1
see also (Goosse et al., 2010a) for the most recent LOVECLIM version 1.2.
estimates of the neglected ageostrophic terms (Lim et al., 1991) in
order to close the equations at the equator. The coupling between
dynamics and thermodynamics is done via a linear balance equation. The sea-ice-ocean component of LOVECLIM, CLIO (Campin and
Goosse, 1999; Goosse et al., 1999; Goosse and Fichefet, 1999)
consists of a free-surface primitive equation model with 3 3
resolution, coupled to a thermodynamic-dynamic sea-ice model.
Coupling between atmosphere and ocean is done via the exchange
of freshwater, momentum and heat fluxes. The terrestrial vegetation module of LOVECLIM, VECODE, (Brovkin et al., 1997) simulates
the dynamical vegetation changes and the terrestrial carbon cycle
in response to climatic conditions.
LOCH is a 3-dimensional global model of the oceanic carbon
cycle (Mouchet and Francois, 1996; Menviel et al., 2008b; Goosse
et al., 2010a). The prognostic state variables considered in the
model are dissolved inorganic carbon (DIC), total alkalinity, phosphates (PO34), organic products, oxygen and silica. LOCH is fully
coupled to CLIO, with the same time step. In addition to their
biogeochemical transformations tracers in LOCH experience the
circulation field and diffusion predicted by CLIO. LOCH computes
the export production from the fate of a phytoplankton pool in the
euphotic zone (0e120 m). The phytoplankton growth depends on
the availability of nutrients (PO34) and light, with a weak
temperature dependence. A grazing process together with natural
mortality limit the primary producers biomass and provide the
source term for the organic matter sinking to depth. The other
processes described in the model include remineralization,
carbonate precipitation and dissolution as well as opal production.
LOCH does not include sedimentary processes but nevertheless
takes into account carbonate compensation mechanisms. The
atmospheric CO2 content is updated for each ocean time step. For
our experiments, however, the atmospheric CO2 was prescribed.
The coupled model has been used extensively to study different
aspects of Late Quaternary climate (Timm and Timmermann, 2007;
Menviel et al., 2008b, 2008a; Timm et al., 2008, 2010; Timmermann
et al., 2009a).
2.2. Experimental setup
In this study the goal is to simulate the sequence of millennialscale events superimposed on the slowly-varying orbital-scale
background conditions during the last deglaciation. For this we
perform two transient experiments starting from the background
conditions of the Last Glacial Maximum (LGM) (Timm and
Timmermann, 2007; Menviel et al., 2008b). The model is then
forced for the period 21 ka to 10 ka B.P. with the time-varying
evolution of solar insolation (Berger, 1978), icesheet topography
(updated every 100 years) (Peltier, 1994), high latitude albedo and
atmospheric CO2. The changes in atmospheric CO2 are obtained
from the EPICA Dome C ice core, Antarctica (Monnin et al., 2001) on
the EDC3 timescale (Lüthi et al., 2008). Other greenhouse gases
(CH4 and N2O) are kept fixed at LGM values.
In experiment DGNS, a series of freshwater pulses is applied in
both the North Atlantic and the Southern Ocean in order to reproduce millennial-scale events (Fig. 1d). The timing and amplitude of
the freshwater pulses were set so as to provide a good match with
the 231Pa/230Th data from the North Atlantic (McManus et al., 2004)
and temperature variations in Greenland (Alley, 2000) on the GISP2
timescale, which is within dating uncertainties of the GICC05
timescale (Rasmussen et al., 2006). Even though a significant
amount of freshwater was probably discharged into the Gulf of
Mexico during the last deglaciation (Aharon, 2003; Flower et al.,
2004; Tarasov and Peltier, 2006; Sionneau et al., 2010), with
potential implications for the AMOC (Roche et al., 2007), we do not
add freshwater into this region. Given the uncertainties associated
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L. Menviel et al. / Quaternary Science Reviews 30 (2011) 1155e1172
1159
Fig. 3. Map showing the locations of the paleo-proxy records used to compare model results. Dark blue stars represent paleo-proxy records of oceanic circulation (McManus et al., 2004;
Piotrowski et al., 2004; Gherardi et al., 2005; Okazaki et al., 2010; Thornalley et al., 2011). Light blue stars represent terrestrial and marine temperature records (Petit et al., 1999; Alley,
2000; Bard et al., 2000; Lea et al., 2003; Pahnke et al., 2003; Nielsen, 2004; Kaiser et al., 2005; Samson et al., 2005; Calvo et al., 2007; Heiri et al., 2007; Jouzel et al., 2007; Martrat et al.,
2007; Barker et al., 2009; Stenni et al., 2011). Orange stars show the location of paleo-precipitation records (Thompson et al.,1995; Peterson et al., 2000; Baker et al., 2001a; 2001b; Wang
et al., 2001; Holmgren et al., 2003; Ivanochko et al., 2005; Wang et al., 2007; Weldeab et al., 2007). Red stars represent biogeochemical paleo-proxy records (Pride et al.,1999; Emmer and
Thunell, 2000; Ivanochko and Pedersen, 2004; Sachs and Anderson, 2005; Anderson et al., 2009). (For interpretation of the references to colour in this figure legend, the reader is
referred to the web version of this article.)
with the detailed freshwater routing during the Last Glacial Termination we choose a more idealized approach. To mimic H1, Older
Dryas and Younger Dryas in our model, freshwater is released into
a broad region in the North Atlantic centered between 55 We10 W,
50 Ne65 N and the Arctic Ocean, respectively. The amplitude of the
millennial-scale freshwater forcing amounts to 0.2 Sv during the
period 18 ka to 17.4 ka B.P. and 0.25 Sv between 17.4 ka B.P and
15.6 ka B.P. To specifically represent the Older Dryas we apply
a freshwater pulse in the North Atlantic for the period 14.4 ka B.P. to
14 ka B.P. with a maximum input of 0.25 Sv between 14.2 and 14 ka
B.P. The exact shape of the freshwater pulse is depicted in Fig. 1d. A
background freshwater flux was also added in the North Atlantic
between 14 and 13 ka B.P. to represent the melting northern hemisphere ice sheets. To mimic the Younger Dryas, 0.25 Sv is applied in
the Arctic Ocean (175 We95 W, 67 Ne83 N) between 13 and
12.2 ka B.P. Finally, to represent the ACR, an additional meltwater
pulse is prescribed for the area enclosing the Ross and Weddell Seas
(163 E11 E, 70 S-80 S). The magnitude of this pulse is 0.15 Sv for
the period 14.4 ka to 13.4 ka B.P., after which the pulse linearly
decreases to 0 over 1000 years (until 12.4 ka B.P.).
To better evaluate the impact of Southern hemispheric meltwater pulses, another experiment similar to DGNS (DGN) was conducted. In DGN, no meltwater pulse is added in the Southern Ocean
between 14.4 ka and 12.4 ka B.P2.
2
Unlike some of the experiments described in Timm and Timmermann (2007),
experiments DGNS and DGN do not employ any orbital acceleration (Lorenz and
Lohmann, 2004).
The results of experiment DGNS are compared to a variety of
paleo-proxy records (Fig. 3 and Table 1), which represent variations
in oceanic circulation, temperature, precipitation and biogeochemical variables in different regions of the world. To obtain
a more representative comparison of model outputs with paleoproxy records, when possible, composite time series were
generated. These composite time series were produced by first
interpolating each paleo-proxy data onto an equidistant grid in
time using bicubic splines and then averaging the different paleoproxy records. In the model-paleo data comparison conducted
here, we allowed for some dating uncertainties in the proxy
records. Optimal match between the local proxy time series and the
simulated physical variable was sometimes achieved by shifting the
proxy time series by a few hundred years (black arrows in time
series figures, Table 2). The shift that we had to apply to obtain
a good match between model solution and paleo data was smaller
than the respective published absolute dating uncertainties. The
time series shifting depends on the timing of the freshwater forcing
sequences prescribed in the transient simulations.
3. The deglacial history in the northern hemisphere
3.1. Heinrich event 1
During H1, and as a result of the chosen freshwater forcing the
simulated AMOC collapses (Fig. 1a), in agreement with the 231Pa/230Th
data from the Bermuda rise core OCE326-GGC5 (McManus et al.,
2004). Fig. 1a also displays the 231Pa/230Th data from the Iberian
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Table 1
paleo-proxy records used to compare the model outputs of this study.
Location
Core
Lat
Long
Elev (m)
Proxy
Reference
Greenland
Greenland
South Iceland rise
South Iceland rise
South Iceland rise
The Netherlands
Iberian margin
Iberian margin
Iberian margin
Iberian margin
Iberian margin
North Atlantic
Cariaco Basin
Cariaco Basin
Arabian Sea
Gulf of Guinea
China
North Pacific
California basin
California basin
California basin
Peruvian Altiplano
Bolivian Altiplano
Bolivian Altiplano
Brazil
South Africa
South Australia
New Zealand
Chilean margin
Southeast Atlantic Ocean
Southeast Atlantic Ocean
New Zealand
New Zealand
Southern Ocean
Southern Ocean
Antarctica
Antarctica
Antarctica
Antarctica
GISP2
GISP2
RAPiD-10-1P
RAPiD-15-4P
RAPiD-17-5P
Hijkermeer
SU81-18
SU81-18
MD95-2042
MD01-2444
ODP-161-977A
OCE326-GGC5
ODP 1002
PL07-39PC
905
MD03-2707
Hulu Cave
72.6 N
72.6 N
62.6 N
62.2 N
61.3 N
52.5 N
37.8 N
37.8 N
37.5 N
37.8 N
36.1 N
33.42 N
10.7 N
10.7 N
10.46 N
2.5 N
32.5 N
32e55 N
27 N
34.17 N
34.17 N
9S
16e17.5 S
20.14 S
27.2 S
24 S
36.44 S
36.55 S
41 S
41.07 S
41.1 S
45.5 S
45.5 S
53.2 S
53.2 S
72.8 S
75.1 S
75.1 S
78.27 S
38.5 W
38.5 W
17.4 W
17.1 W
19.3 W
6.3 E
10.1 W
10.1 W
10.1 W
10.7 W
1.6 W
57.35 W
65.2 W
65 W
51.57 E
9.4 E
119.2 E
133e168 E
112 W
120.02 W
120.02 W
77.5 W
68.5e70 W
67.3 W
49.2 W
29.2 E
136.33 E
177.26 E
74.5 W
9E
7.8 E
174.9 E
174.9 E
5.1 E
5.1 E
159.2 E
123.4 E
123.4 E
106.5 E
3207
3207
1237
2133
2303
14
3135
3135
3146
3140
1924
4550
893
790
1586
1295
100
Ice core d18O, borehole
Ice core CH4
Ventilation age
Ventilation age
Ventilation age
Chironomids
UK’37
231
Pa/230Th
UK’37
UK’37
UK’37
231
Pa/230Th
550 nm reflectance
Foram Mg/Ca
d15N
Foram Ba/Ca
Speleothem d18O
Ventilation age
d15N
d15N
Mo/Al
Ice core d18O
% freshwater diatom
g nat. radiation
Speleothem d18O
Speleothem d18O
UK’37
Modern analog
UK’37
Nd
Foram. Mg/Ca
Alkenone content
UK’37
Modern analog
Opal flux
Ice core d18O
Ice core dD
CO2
Ice core dD
(Alley, 2000)
(Brook et al., 1996)
(Thornalley et al., 2011)
(Thornalley et al., 2011)
(Thornalley et al., 2011)
(Heiri et al., 2007)
(Bard et al., 2000)
(Gherardi et al., 2005)
(Bard, 2002)
(Martrat et al., 2007)
(Martrat et al., 2004)
(McManus et al., 2004)
(Peterson et al., 2000)
(Lea et al., 2003)
(Ivanochko et al., 2005)
(Weldeab et al., 2007)
(Wang et al., 2001)
(Okazaki et al., 2010)
(Pride et al., 1999)
(Emmer and Thunell, 2000)
(Ivanochko and Pedersen, 2004)
(Thompson et al., 1995)
(Baker et al., 2001a)
(Baker et al., 2001b)
(Wang et al., 2007)
(Holmgren et al., 2003)
(Calvo et al., 2007)
(Samson et al., 2005)
(Kaiser et al., 2005)
(Piotrowski et al., 2004)
(Barker et al., 2009)
(Sachs and Anderson, 2005)
(Pahnke et al., 2003)
(Nielsen, 2004)
(Anderson et al., 2009)
(Stenni et al., 2011)
(Jouzel et al., 2007)
(Monnin et al., 2001)
(Petit et al., 1999)
JPC-56
ODP893-A
ODP893-A
Huascaran
Lake Titicaca
Salar de Uyuni
Botuverá Cave
Cold Air Cave
MD03-2611
H214
ODP 1233
RC11-83
TN057-21
MD97-2120
MD97-2120
TN057-13PC
TN057-13PC
TALDICE
EPICA Dome C
EPICA Dome C
Vostok
margin core SU81-18 (Gherardi et al., 2005) and the ventilation age as
inferred from South Iceland rise sediment cores (Thornalley et al.,
2011). Even if all the records show a weakened AMOC at some time
during H1, there are some substantial differences among them, which
will further be discussed in section 5. The AMOC weakening reduces
the poleward Atlantic heat transport at 40 N by 0.5 PW (not shown),
which leads to an enhanced Northern Hemispheric cooling relative to
the LGM (Figs. 2 and 4 left), in accordance with temperature paleoproxy data from Greenland (Alley, 2000) (Fig. 2a), the North Atlantic
(Bard et al., 2000; Calvo et al., 2001; Waelbroeck et al., 2001; Bard,
2002; Martrat et al., 2004, 2007) (Fig. 2c), the tropical North Atlantic
(Lea et al., 2003; Niedermeyer et al., 2009) (Fig. 2d), the Mediterranean
Sea (Cacho et al., 1999, 2001), Europe (Ivy-Ochs et al., 2006), Asia
(Wang et al., 2002; Evans et al., 2003) and Africa (Powers et al., 2005).
818
576.5
576.5
6048
3810
3653
250
e
2420
2045
838
4718
4981
1210
1210
2848
2848
2315
3240
3240
3500
The enhanced meridional temperature gradient in the North Atlantic
induces a strengthening of the northeasterly trades and a southward
shift of the Intertropical Convergence Zone (ITCZ) (Timmermann et al.,
2005) (Fig. 4, middle). In agreement with paleo-proxies, drier conditions prevailed during H1 in Europe (Peyron et al., 2005), the northern
part of Africa (Peck et al., 2004; Williams et al., 2006; Lamb et al., 2007;
Niedermeyer et al., 2009), the Middle East (Bar-Matthews et al., 2003;
Vaks et al., 2003; Shakun et al., 2007; Fleitmann et al., 2009) and the
northern part of South America. As shown in Fig. 5a, we simulate dry
conditions in the Cariaco basin as soon as the AMOC weakens. The
reflectance record from the Cariaco sediments (Peterson et al., 2000)
indicates relatively dry conditions between 18 and 16.2 ka B.P.,
however the drying becomes more intense at 16.2 ka B.P. The reason
behind this feature remains unexplained. In the Gulf of Guinea, the
Table 2
Temporal shift applied to paleo-proxy records used to compare the model outputs of this study. A positive (negative) shift means that the paleo data had to be shifted forward
(backward) in time to match the model simulation.
Location
Cores
Proxy
Reference
Temporal shift applied (yrs)
North Atlantic
North Pacific
Bolivian Altiplano
Bolivian Altiplano
Southeast Atlantic Ocean
New Zealand
Southern Ocean
Antarctica
Antarctica
OCE326-GGC5
231
(McManus et al., 2004)
(Okazaki et al., 2010)
(Baker et al., 2001a)
(Baker et al., 2001b)
(Piotrowski et al., 2004)
(Sachs and Anderson, 2005)
(Anderson et al., 2009)
(Jouzel et al., 2007)
(Petit et al., 1999)
() 200
() 600
400
400
500
() 300
() 200
() 600
() 600
Lake Titicaca
Salar de Uyuni
RC11-83
MD97-2120
TN057-13PC
EPICA Dome C
Vostok
Pa/230Th
Ventilation age
% freshwater diatom
g nat. radiation
Nd
Alkenone content
Opal flux
Ice core dD
Ice core dD
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L. Menviel et al. / Quaternary Science Reviews 30 (2011) 1155e1172
1161
Fig. 4. (left) Annual SST anomalies ( C, shaded) and current anomalies (m/s, vectors) averaged over 0e200 m for H1eLGM conditions (16e18 ka B.P.), (middle) Annual precipitation
anomalies (cm/yr, shaded) and annual wind stress anomalies (Pa, vector), (right) Annual export production anomalies (gC/m2/yr) for H1eLGM and overlaid 0.1 m annual sea-ice
contour for H1.
Fig. 5. (a) Time series of JJA precipitation anomalies (cm/yr) averaged over the Cariaco basin (55 W-52 W, 10 N-14 N) for experiment DGNS (green) compared with the
reflectance (%, black) measured in a marine sediment core from the Cariaco basin (ODP 1002) (Peterson et al., 2000). Lower reflectance has been associated with increased marine
productivity due to greater riverine input. (b) Time series of JJA precipitation anomalies (cm/yr) averaged over western Equatorial Africa (15 W-15 E, 4 N-20 N) for experiment
DGNS (green) compared with SSS reconstruction (black) from the Gulf of Guinea core MD03-2707 (Weldeab et al., 2007). (c) Time series of JJA/DJF precipitation averaged over China
(114 E124 E, 28 N-35 N) for experiment DGNS (green) compared to d18O (permil, black) as recorded in stalagmites of the Hulu cave (China) (Wang et al., 2001). (d) Time series of
JJA precipitation anomalies over the Arabian Sea (45 E65 E, 5 N-15 N) for experiment DGNS (green) compared to d15N (permil, black) as recorded in marine sediment core 905
from the Arabian Sea (Ivanochko et al., 2005). Anomalies are with respect to pre-industrial time. Time series are 51 year running means for the model outputs. (For interpretation of
the references to colour in this figure legend, the reader is referred to the web version of this article.)
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L. Menviel et al. / Quaternary Science Reviews 30 (2011) 1155e1172
southward ITCZ shift leads to drier conditions in agreement with Ba/Ca
data indicating greater surface salinity during H1 in that area (Weldeab
et al., 2007). The AMOC shut down also leads to a reduced ratio of
summer/winter precipitation over China in agreement with d18O
records from Chinese caves (Wang et al., 2001; Dykoski et al., 2005)
(Fig. 5c). Furthermore, our model solution supports the notion of
a reduced summer monsoon over the Arabian Sea (Ivanochko et al.,
2005) (Fig. 5d).
As a result of the simulated southward shift of the ITCZ (Fig. 4,
middle), the reduced Atlantic-Pacific freshwater export (not
shown) and the reduced Asian summer monsoon (Fig. 5c), North
Pacific salinity increases at the onset of H1, as described in Okazaki
et al. (2010). North Pacific waters become dense enough to generate
North Pacific Deep Water (NPDW) formation in the Bering Sea and
the Gulf of Alaska. Subsequently a deep Pacific Meridional Overturning Cell (PMOC) establishes. Maximum meridional transports
in this cell attain values of up to 20 Sv. Southward spreading of
well-ventilated NPDW is in accordance with a compilation of
radiocarbon data obtained from marine sediment cores that
suggest a drop of eastern North Pacific ventilation ages by w1000
years (Fig. 1b) Okazaki et al. (2010). As can be seen in Fig. 4 left, the
formation of NPDW leads to a reorganization of surface currents in
the Indo-Pacific: the Indonesian Throughflow transport decreases
(w-50%) and the warm and saline water of the subtropical Pacific
are diverted into the northern North Pacific, supplying the sinking
branch of the PMOC (Chikamoto et al., 2010; Menviel et al., 2010b;
Okazaki et al., 2010). A slowdown of the Indonesian Throughflow
during Heinrich events has also been inferred from a surface and
upper thermocline temperature record in the Timor Sea (Zuraida
et al., 2009), providing support to our analysis. The formation of
NPDW leads to an increase in poleward heat transport in the North
Pacific of 0.5 PW at 40 N. In experiment DGNS, the enhanced
transport of warm equatorial waters to the North Pacific combined
with greater insolation and atmospheric CO2 lead to a slight
warming of the ocean surface in the North Pacific during H1
(16e18 ka B.P.) (Fig. 4, left). In agreement with SST paleo-proxies,
a warming is observed in the northeastern Pacific where warm
waters converge (Sarnthein et al., 2006), while a relatively small
area around Japan cools (Harada et al., 2006, 2010) (Fig. 4, left). At
the beginning of H1, the ITCZ shifts southward in the Pacific,
however as NPDW formation enhances, the ITCZ shifts back to its
initial position in the Pacific basin (Fig. 4, middle).
In the North Pacific, the formation of NPDW leads to an increase
in the dissolved oxygen content (by up to 3 ml/L) in the upper
2500 m, in agreement with benthic foraminiferal assemblages in
marine sediment cores from the Northwest Pacific (Shibahara et al.,
2007). The well-oxygenated waters flow equatorwards from the
North Pacific. The main branch flows as a deep western boundary
current but well-ventilated waters are also advected along the
California basin. The dissolved oxygen content along the coast of
California increases by about 50% down to about 2000 m depth.
This is in qualitative agreement with redox sensitive metal content,
d15N and total organic carbon records from marine sediment core
ODP 893A (Kennett and Ingram, 1995; Emmer and Thunell, 2000;
Ivanochko and Pedersen, 2004) as well as with d15N record from
marine sediment core JPC-56 (Pride et al., 1999) (Fig. 1c). Despite
the fact that changes in d15N can have numerous causes (Somes
et al., 2010), the redox sensitive metal content and organic carbon
records from the California basin support our hypothesis of greater
ventilation during H1.
Due to the colder conditions and increased stratification,
simulated marine export production is low in an area spanning the
North Atlantic, the Atlantic coast of North and Central America as
well as the Caribbean Sea, in agreement with paleo-proxy reconstructions (Peterson et al., 2000; Vink et al., 2001) (Fig. 4, right).
Due to the stronger upwelling, marine export production significantly increases (þ25%) in the Equatorial Atlantic (Northwest
African coast and Brazilian coast) in agreement with paleo-proxy
records from the Brazilian coast (Jennerjahn et al., 2004).
3.2. Bølling warming, Older Dryas and Intra-Allerød cold period
At the end of H1 (w15.2 ka B.P.), the simulated AMOC recovers
quickly regaining a strength of about 27 Sv around 14.6 ka B.P
(Fig. 1a). As can be seen in Figs. 2 and 6 (left), the temperatures in
Greenland and in the North Atlantic increase abruptly during the
recovery phase of the AMOC (15e14.6 ka B.P.), until reaching
a maximum during the Bølling Warm Period (BWP). As already
discussed by Liu et al. (2009), the rapid transition to the BWP can be
explained by a combination of a rapid resumption of the AMOC
following H1, an overshooting effect of the AMOC, as well as by the
contemporaneous increase in atmospheric CO2 concentrations
(Köhler et al., 2010). The amplitude of the simulated increase in
spring temperature over Greenland between 15.2 ka B.P. and 14.5 ka
B.P. amounts to about 12 C, which is similar to the 13 C temperature difference estimated from GISP2 ice core for the same time
period (Alley, 2000) (Fig. 2a). Off the Iberian margin the simulated
SST and alkenone-based SST (Bard et al., 2000; Bard, 2002; Martrat
et al., 2004, 2007) both increase between H1 and BWP by about
3e4 C (Fig. 2c). In the Cariaco basin however, the simulated 1.5 C
SST increase during the winter season is only half of that reconstructed from Mg/Ca data in core PL07-39PC (Lea et al., 2003)
(Fig. 2d). The resumption of the AMOC leads to a reorganization of
surface ocean currents and to the decline in the formation of NPDW.
Fig. 6. (left) Annual SST anomalies ( C, shaded) and current anomalies (m/s, vectors) averaged over 0e200 m for Bølling warmingeH1 conditions (14.6e16 ka B.P.), (middle) Annual
precipitation anomalies (cm/yr, shaded) and annual wind stress anomalies (Pa, vector), (right) Annual export production anomalies (gC/m2/yr) for BWPeH1 and overlaid 0.1 m
annual sea-ice contour for BWP.
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L. Menviel et al. / Quaternary Science Reviews 30 (2011) 1155e1172
Such an intense and rapid warming in the North Atlantic at
about 15 ka B.P. is likely to have consequences on the massive ice
sheets that are still present over North America and northern
Europe. During this period, melting of the Laurentide and Fennoscandian ice sheets accelerated (Peltier, 1994; Gregoire, 2010;
Thornalley et al., 2010) thus adding freshwater into the North
Atlantic. In response to this freshwater perturbation the simulated
AMOC weakens from 27 Sv at the peak of the Bølling warming
down to about 12 Sv during the Older Dryas (w14 ka B.P.) and levels
off at about 18 Sv during the Allerød (between 13.8 and 13 ka B.P.)
(Fig. 1a). These variations in simulated AMOC strength are in overall
agreement with the ventilation age history recorded in the South
Iceland rise (Thornalley et al., 2011). The significant freshwaterinduced reduction in AMOC strength during the Older Dryas is not
observed in the 231Pa/230Th records (McManus et al., 2004;
Gherardi et al., 2005). Higher resolution 231Pa/230Th records for
the period 14.2 and 13.5 ka B.P would be needed to test this
assertion. Due to the intermittent AMOC weakening at 14 ka
B.P., modeled air temperatures over Greenland drop by about 6 C,
summer air temperature decreases by 1.8 C in the Netherlands, SST
decrease by about 1.2 C off the Iberian margin and by about 0.7 C
in the Cariaco basin (Figs. 2 and 7, left). This simulated cooling is in
very good agreement with temperature estimates from Greenland
ice core (Alley, 2000) and chironomid-inferred July temperatures
from the Netherlands (Heiri et al., 2007). However the cooling in
the Cariaco basin SST reconstruction (Lea et al., 2003) is lower than
in the model results and no significant cooling is observed in
marine sediment cores off the Iberian margin (Bard et al., 2000;
Bard, 2002; Martrat et al., 2004, 2007).
Over northern North America, the waning of the icesheet and
the resulting changes in height induce a local surface warming due
to the lapse rate effect, even during the Older Dryas. At about
13.8 ka B.P. the simulated temperatures in the North and Equatorial
Atlantic recover to levels equivalent to the Bølling warming, in
agreement with SST proxies from that region.
Greenland ice cores also record another millennial-scale cold
event, known as the Intra-Allerød cold period (IACP, 13.5e13 ka B.P.,
Fig. 2a) (Stuiver et al., 1995), which is not reproduced in our transient modeling experiments. This millennial-scale cold event was
also observed in the Norwegian Sea (Lehman and Keigwin, 1992),
on the South Iceland Rise (Thornalley et al., 2010) as well as in
terrestrial records from Northern Europe (von Grafenstein et al.,
1999; Heiri et al., 2007) (Fig. 2b) and North America (Yu and
Eicher, 2001). In addition ventilation age anomalies indicate
decreased ventilation south of Iceland during that period
(Thornalley et al., 2011). However no significant decrease in North
Atlantic SST was reconstructed for that time period according to
1163
higher resolution SST proxies (Cacho et al., 1999, 2001; Bard et al.,
2000; Calvo et al., 2001; Bard, 2002; Lea et al., 2003; Martrat
et al., 2004, 2007). Given these discrepancies between the proxybased spatial extent of the IACP and the model simulation it thus
remains unresolved whether large-scale freshwater forcing was the
sole cause for these variations.
During the Bølling warm period, the ITCZ swings back to the
north, leading to wetter conditions over Europe (Peyron et al.,
2005), the northern part of South America (Peterson et al., 2000,
North Africa (Peck et al., 2004; Williams et al., 2006; Weldeab et al.,
2007; Niedermeyer et al., 2009), the middle East (Bar-Matthews
et al., 2003; Vaks et al., 2003; Ivanochko et al., 2005; Shakun
et al., 2007; Fleitmann et al., 2009), China (Wang et al., 2001;
Dykoski et al., 2005) and Tibet (Kramer et al., 2010) (Fig. 6,
middle and 5).
Due to the intense warming and retreat of the sea-ice edge in
the North Atlantic the marine export production more than doubles
in this region (Fig. 6, right). It also increases considerably in the
Caribbean Sea (Peterson et al., 2000; Vink et al., 2001). Marine
export production however decreases in the North Pacific (5%)
due to the reduction of the NPDW formation and in the Equatorial
Atlantic (25%) due to the reduction in the strength of the northeasterly trades (Jennerjahn et al., 2004).
3.3. Younger Dryas
In two recent studies (Tarasov and Peltier, 2005; Murton et al.,
2010) it was suggested that a glacial meltwater pulse into the
Arctic Ocean caused the YD cold period. The shut down of the
AMOC at 13 ka B.P. (Fig. 1a) leads to a cooling of about 10 C over
Greenland, 2 C in the Netherlands, 3 C off the Iberian margin and
1 C in the Cariaco basin (Fig. 2). The simulated relative strength of
the overturning in the North Atlantic is weaker than relative
changes recorded by the 231Pa/230Th record of the Bermuda rise and
Iberian margin sediment cores (McManus et al., 2004; Gherardi
et al., 2005). Ventilation age anomalies from the South Iceland
rise indicate an interruption of the ventilation in the second part of
the YD (Thornalley et al., 2011). Over Greenland, the amplitude of
the simulated cooling during the YD is similar to the one estimated
from the GISP2 borehole (Cuffey and Clow, 1997; Alley, 2000), but
the simulated temperature at 13 ka B.P. (before the YD) is about 5 C
warmer than the one estimated in GISP2 (Fig. 2a). This 5 C offset
from the observations lasts until the end of the experiment DGNS.
Part of this temperature discrepancy between the model and the
GISP2 record could be explained by an elevation change at GISP2.
Using the relationship between height and d18O (Vinther et al.,
2009), a 200 m increase in the GISP2 icesheet would lead to a 1.2
Fig. 7. (left) Annual SST anomalies ( C, shaded) and current anomalies (m/s, vectors) averaged over 0e200 m for ACReBølling warming conditions (14e14.6 ka B.P.), (middle)
Annual precipitation anomalies (cm/yr, shaded) and annual wind stress anomalies (Pa, vector), (right) Annual export production anomalies (gC/m2/yr) for ACReBWP and overlaid
0.1 m annual sea-ice contour for ACR.
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L. Menviel et al. / Quaternary Science Reviews 30 (2011) 1155e1172
permil decrease in d18O. Assuming a slope of 2.8 C per 1 permil for
the deglaciation, the recorded temperatures would be w3.3 C
lower. On top of that the adiabatic lapse rate (0.98 C/100 m) has to
be taken into account. A 5 C temperature offset could thus be
explained by a 180e200 m increase in the icesheet height at GISP2.
The simulated cooling off the Iberian margin is overestimated
compared to the estimated SST anomaly (Bard et al., 2000; Bard,
2002; Martrat et al., 2004, 2007), while it is in agreement with
a reconstructed 3 C cooling in the Western Mediterranean Sea
(Cacho et al., 1999, 2001). In the Cariaco basin, however, the
simulated cooling is underestimated. The abrupt YD cooling is also
observed in paleorecords from continental Europe (von Grafenstein
et al., 1999; Heiri et al., 2005; Heiri and Millet, 2005; Peyron et al.,
2005; Genty et al., 2006).
Similar to the H1 climate response, the anomalous meridional
SST gradient during the YD in the tropical Atlantic leads to
a southward shift of the ITCZ. In agreement with paleo-proxies, drier
conditions are simulated over the northern part of South America
(Maslin and Burns, 2000; Peterson et al., 2000; Haug et al., 2001),
Europe (Goslar et al., 1999; Kerschner et al., 2000; Peyron et al.,
2005), Northern Africa (Peck et al., 2004; Williams et al., 2006;
Lamb et al., 2007; Weldeab et al., 2007; Niedermeyer et al., 2009)
and the Middle East (Bar-Matthews et al., 2003; Vaks et al., 2003;
Ivanochko et al., 2005; Shakun et al., 2007; Fleitmann et al., 2009).
Fig. 8. (a) Time series of the maximum overturning strength in the bottom cell of the Southern Ocean (Sv) for experiment DGNS (green) and DGN (cyan) compared to eNd (black) as
recorded in marine sediment core RC11-83 (Piotrowski et al., 2004). Lower eNd values indicate the predominance of NADW at the site of core RC11-83 while higher values indicate
the predominance of AABW. The time series of (Piotrowski et al., 2004) has been shifted by 500 years forward. (b) Time series of austral spring (SON) air temperature anomalies ( C)
averaged over 70 S-90 S for experiments DGNS (green) and DGN (cyan) compared to temperature reconstructions ( C) from EPICA Dome C ice core (black stars) on the EDC(3) time
scale (Jouzel et al., 2007) and from Vostok ice core (black thin line) (Petit et al., 1999) as well as compared to the d18O record from Talos Dome ice core (gray) (Stenni et al., 2011). The
time axis of the EPICA Dome C and Vostok temperature reconstructions were shifted backward by 600 years to better fit the model results and the northern hemispheric paleoproxies. Simulated austral spring temperatures were used as suggested by the study of (Timmermann et al., 2009b). (c) Time series of annual SST anomalies ( C) averaged over the
Pacific and Atlantic part of the Southern Ocean (170 E10 E, 40 S-55 S) for experiments DGNS (green) and DGN (cyan) compared to a composite of SST reconstructions ( C, black)
from marine sediment cores MD97-2120 (Pahnke et al., 2003), ODP1233 (Kaiser et al., 2005), TN057-13PC4 (Anderson et al., 2009; Nielsen, 2004) and TN057-21 (Barker et al., 2009).
(d) Time series of annual SST anomalies ( C) averaged over the Southwest Pacific (130 E180 E, 32 S-42 S) for experiment DGNS (green) compared to a composite of SST reconstructions ( C, black) from marine sediment cores MD03-2611 (Calvo et al., 2007) and H214 (Samson et al., 2005). Anomalies are with respect to pre-industrial time. Time series are
21 year running means for the model outputs. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)
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L. Menviel et al. / Quaternary Science Reviews 30 (2011) 1155e1172
Reduced summer precipitation is also simulated over China (Wang
et al., 2001; Dykoski et al., 2005) (Fig. 5).
The weakening of the AMOC during the YD leads to the
formation of NPDW. As a result, younger ventilation ages are
recorded in marine sediment cores from the North Pacific (Fig. 1b)
(Okazaki et al., 2010) and dissolved oxygen content increases
along the coast of California (Kennett and Ingram, 1995; Pride
et al., 1999; Emmer and Thunell, 2000; Ivanochko and Pedersen,
2004) (Fig. 1c).
4. The deglaciation history in the Southern Hemisphere
4.1. Heinrich event 1
As a result of the simulated AMOC shut down during H1, the
AABW cell strengthens, attaining values of up to 12 Sv between 17.6
1165
and 14.8 ka B.P. (Fig. 8a). This intensification of AABW formation is
in qualitative agreement with Neodynium data from the South
Atlantic core RC11-83 (Piotrowski et al., 2004). Lower eNd values
indicate the predominance of NADW at the site of core RC11-83
while higher values indicate the predominance of AABW. Accompanying the NADW decrease and the AABW increase, the model
simulates an intensification of the poleward heat transport at 30 S
by about 0.6PW. As depicted in Fig. 8, temperatures at high
southern latitudes increase almost linearly from 17.6 to 14.8 ka B.P.
This warming is a combination of different factors: rising atmospheric CO2 concentrations (Timmermann et al., 2009b), increase of
austral spring mean insolation changes affecting the sea-ice extent
around Antarctica, and the bipolar seesaw effect leading to
a warming of the Southern Hemisphere. Over Antarctica, the
simulated air temperature during austral spring follows quite
closely the temperature derived from EPICA Dome C ice core dD
Fig. 9. (a) Time series of annual precipitation anomalies (cm/yr) averaged over Bolivia (71 W-62 W, 15 S-25 S) for experiment DGNS (green) compared to % of freshwater diatoms
(black x) in Lake Titicaca (Bolivia) (Baker et al., 2001b) and to a natural gamma ray profile (c.p.s., gray circles) from the Salar de Uyuni (Bolivia) (Baker et al., 2001a). The time axis of
the Lake Titicaca and the Salar de Uyuni records were shifted forward by 400 years. (b) Time series of DJF precipitation anomalies (cm/yr) averaged over Peru (85 W-70 W, 5 S
18
15 S) for experiment DGNS (green) compared to d O record (permil, black) from Huascaran ice core (Peru) (Thompson et al., 1995). (c) Time series of DJF/JJA precipitation averaged
over Brazil (60 W-44 W, 20 S-32 S) for experiment DGNS (green) compared to d18O record (permil, black) from stalagmites of Botuverá cave (Brazil) (Wang et al., 2007). (d) Time
series of JJA/DJF precipitation anomalies averaged over South Africa (25 E35 E, 20 S-30 S) for experiment DGNS (green) compared to d18O record (permil, black) from stalagmites of
the Makapansgat Valley (South Africa) (Holmgren et al., 2003). Anomalies are with respect to pre-industrial time. Time series are 51 year running means for the model outputs. (For
interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)
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L. Menviel et al. / Quaternary Science Reviews 30 (2011) 1155e1172
2
Opal production (mol/m /yr)
a
17
16
4.2. Antarctic Cold Reversal
Simulated surface temperatures at high southern latitudes reach
a local maximum around 14.8 ka B.P (Fig. 8bec). The rapid
strengthening of the AMOC towards the end of H1 leads to
a weakening of AABW formation starting 14.8 ka B.P. (Fig. 8a). The
warmer conditions in the Southern Ocean as well as higher sea
level around 14.6 ka B.P. (Bindschadler et al., 2003; Peltier, 2005),
may have helped to destabilize the Antarctic ice shelves and ice
sheets (Flückiger et al., 2006), thus adding freshwater to the
Southern Ocean. Proxy-evidence for a meltwater pulse originating
from Antarctica at around 14.6 ka B.P. was found in several marine
sediment cores of the Southern Ocean and was discussed in section
1. Adding freshwater to the Southern Ocean in our simulation DGNS
further weakens the formation of AABW and slows down the
Southern Ocean overturning cell to 5 Sv until about 13 ka B P, in
qualitative agreement with Neodynium data from the South
Atlantic (Piotrowski et al., 2004) (Fig. 8a). In contrast, when no
freshwater is added to the Southern Ocean (experiment DGN), the
AABW stays moderately strong (w12 Sv). The latter scenario does
not agree well with the Neodymium data.
Reduced AABW formation in experiment DGNS leads to
a decrease in southward heat transport at 30 S of about 0.2 PW. As
a result, temperatures decrease at high and mid Southern latitudes
(Fig. 7, left). Austral spring temperature averaged over Antarctica
are about 0.8 C lower (Fig. 8b), which compares reasonably well
with the cooling recorded in the dD data from EPICA Dome C (Jouzel
14
13
12
Southern Ocean
0.15
6
ACR
0.1
4
2
H1
0.05
17
36
b
17
16
16
15
14
15
14
13
13
12
12
2
Export production (gC/m /yr)
15
Due to the warmer conditions during H1 and retreating sea-ice
in the Southern Ocean, simulated marine export production
increases in these areas by about 15% (Fig. 10 a and 4, right), in close
correspondence with the data of Anderson et al. (2009) and Sachs
and Anderson (2005).
Opal flux (g/cm2/kyr)
record (Jouzel et al., 2007). Both show an increase of about 4 C
between 18 and 15 ka B.P. At that time, the derived temperature
from the EPICA Dome C ice core suddenly increases by about 2 C in
200 years (Fig. 8b). This temperature spike is not simulated by our
model. As austral spring insolation is an important pacemaker of
the glacial termination in the Southern Hemisphere (Timmermann
et al., 2009b), simulated austral spring temperatures were used for
the comparison with Antarctic records.
In the Atlantic and Pacific sectors of the Southern Ocean
(Fig. 8c), the simulated annual mean SST increases by 2.7 C
between 18 and 15 ka B.P., which is in agreement with the
reconstructed SST increase from a composite of marine sediment
cores that we obtained by averaging linearly interpolated alkenone SST reconstructions for cores MD97-2120 (Pahnke et al.,
2003), ODP1233 (Kaiser et al., 2005), TN057-13PC4 (Nielsen,
2004; Anderson et al., 2009) and TN057-21 (Barker et al., 2009).
At the same time period, the simulated SST increases by 1.9 C in
the South Pacific (32 S-42 S, Fig. 8d). A composite of SST reconstructions from marine sediment cores MD03-2611 (Calvo et al.,
2007) and H214 (Samson et al., 2005) indicates a larger increase
of the order of 3 C, but the overall timing matches well with the
transient simulation.
The cooling of the North Atlantic during H1 in combination with
the warming of the Southern Hemisphere lead to a southward shift
of the ITCZ and generally wetter conditions over the Southern
Hemisphere (Fig. 4, middle), in agreement with paleoproxies from
the Bolivian (Baker et al., 2001a, 2001b; Blard et al., 2009) and
Peruvian (Thompson et al., 1995) Altiplano (Fig. 9aeb) as well as
with d18O records from South African stalagmites (Holmgren et al.,
2003) (Fig. 9d). The South American Summer Monsoon is also quite
intense during H1, as suggested by the d18O recorded in stalagmites
from Botuverá cave (Brazil) (Cruz et al., 2005; Wang et al., 2007)
and our modeling results (Fig. 9c).
Southern Ocean
600
32
400
28
Alkenone content (ng/g)
1166
200
17
16
15
14
13
12
Time (ka BP)
Fig. 10. (a) Time series of opal production (mol/m2/yr) in the Southern Ocean (52 S-64 S) for experiments DGNS (green) and DGN (cyan) compared to the opal flux record (g/cm2/kyr,
black) from core TN057-13 (Anderson et al., 2009). The time axis of the opal flux was shifted backward by 200 years. (b) Time series of marine export production (gC/m2/yr) in the
Pacific side of the Southern Ocean (150 E190 E, 42 S-49 S) for experiments DGNS (green) and DGN (cyan) compared to the alkenone content record (ng/g, black) from core MD972120 (Sachs and Anderson, 2005). The alkenone record is shown for the period 18 to 10.4 ka B.P. (For interpretation of the references to colour in this figure legend, the reader is
referred to the web version of this article.)
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L. Menviel et al. / Quaternary Science Reviews 30 (2011) 1155e1172
1167
Fig. 11. Austral spring air temperature anomalies (14e14.8 ka B.P, C) over Antarctica and Southern Ocean for (left) experiment DGNS and (right) experiment DGN.
et al., 2007) and Vostok ice cores (Petit et al., 1999) and qualitatively
with the d18O from Taylor Dome ice core (Steig et al., 1998). While
the EPICA Dome C and Vostok ice core records had to be shifted by
600 years, the d18O record from Talos Dome ice core based on the
new GICC05 timescale fits well with our simulation without any
time shift. Over the Pacific and Atlantic section of the Southern
Ocean (40 S-55 S), the simulated cooling of 1.4 C is in agreement
with a reconstructed 1.2 C cooling in the Southern ocean SST index
derived in this work from various datasets (Pahnke et al., 2003;
Nielsen, 2004; Kaiser et al., 2005; Anderson et al., 2009; Barker
et al., 2009) (Fig. 8c). Finally in the South Pacific (32 S-42 S),
a simulated SST decrease of 0.9 C compares well with a recon
structed SST decrease of 1.2 C (Calvo et al., 2007; Samson et al.,
2005) (Fig. 8d). Alkenone-based SST estimates from marine sedi
ment core GIK 17748-2 (32.45 S,72.02 W) (not shown) also
suggest a similar cooling of 1 C during the ACR (Kim et al., 2002).
A part of the cooling in the Southern Hemisphere prior to the
addition of freshwater is caused by the reduction of AABW formation
at 14.8 ka B.P. When no freshwater is added to the Southern Ocean
(experiment DGN), the cooling amounts to 0.3 C over Antarctica and
0.6 C in the Southern Ocean. The cooling at high Southern latitudes
in experiment DGNS is more pronounced over the ocean than over
Antarctica (Fig. 11). In experiment DGN the cooling is reduced
significantly compared to DGNS, particularly over Antarctica and the
Atlantic and the Eastern Pacific side of the Southern Ocean.
During the ACR the temperature decrease in the Southern
Hemisphere leads to relatively drier conditions over most of the
Southern Hemisphere (Fig. 7, middle). This is in good qualitative
agreement with paleo-proxies that document decreasing precipitation over the Bolivian and Peruvian Altiplano (Thompson et al.,
1995; Baker et al., 2001a; 2001b) and a reduction of summer
rainfall over South Africa (Holmgren et al., 2003) (Fig. 9). Our model
also simulates a weakening of the South American Summer
Monsoon over Brazil during the ACR in agreement with the d18O
record from Botuverá cave (Brazil) (Wang et al., 2007).
Due to the cooling at high southern latitudes around w14
ka B.P., the simulated summer sea-ice edge advances to about
53e57 S in agreement with paleoproxies (Bianchi and Gersonde,
2004) (contour in Fig. 7, right). The extended sea-ice coverage
induces a decrease in export production in the Southern Ocean due
to light limitation. Both the simulated export production and opal
production in the Southern Ocean significantly decrease during the
ACR (Figs. 10 and 7, right), in agreement with paleoproxies from the
Southern Ocean. Marine sediment core TN057-13 (Anderson et al.,
2009) exhibits a decline in opal flux during the ACR, which corresponds to a decrease in diatom-productivity. In addition, a decrease
in alkenone content indicating a decrease in marine productivity
during the ACR is recorded in marine sediment core MD97-2120
(Sachs and Anderson, 2005). When no freshwater pulse is added to
the Southern Ocean (experiment DGN), the simulated anomalies in
marine export and opal production obtained are much smaller and
last for only 1000 years instead of 1700 years for DGNS (Fig. 10).
4.3. Younger Dryas
The shut down of the AMOC and the associated strengthening of
AABW formation lead to an enhanced poleward heat transport into
the Southern Ocean. Surface temperatures thus increase rapidly in
the Southern Hemisphere. In conjunction with an increase of radiative forcing and changes in obliquity the simulated austral spring air
temperature increases by 2.6 C between 14 and 12 ka B.P., which is
somewhat less than for the temperature reconstructions from the
EPICA Dome C and Vostok ice cores (Petit et al., 1999; Jouzel et al.,
2007) (Fig. 8b). Both the simulated and reconstructed SST in the
Southern Ocean increase by 2 C. However the simulated SST
increases more abruptly than the reconstructed one (Fig. 8c). In the
South Pacific, the simulated SST increases by 1.6 C compared to
a reconstructed SST increase of about 3 C (Fig. 8d). In addition, the
simulated SST decreases by 1 C at 11.5 ka B.P., in contrast to the paleo-
Fig. 12. Time series of annual terrestrial primary production anomalies (GtC/yr)
averaged over the latitudinal band 0e30 S (cyan), 0e20 N (red), 45 Ne70 N
(magenta). The green line represents the sum of the cyan, red and magenta lines. The
black line represents the time series of the atmospheric CH4 (ppb) as recorded in GISP2
ice core (Brook et al., 2000). (For interpretation of the references to colour in this figure
legend, the reader is referred to the web version of this article.)
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L. Menviel et al. / Quaternary Science Reviews 30 (2011) 1155e1172
proxy indices used here. Due to the warmer conditions at high
southern latitudes and reduced sea-ice coverage, the export and opal
productions also increase strongly in agreement with paleoproxies
data (Sachs and Anderson, 2005; Anderson et al., 2009) (Fig. 10).
In agreement with paleorecords, wetter conditions prevail over
the Bolivian Altiplano (Fig. 9a) (Baker et al., 2001a; 2001b). Wetter
conditions are also simulated over the Peruvian Altiplano (Fig. 9b),
while the Huascaran d18O records show only a slight increase in
precipitation (Thompson et al., 1995). Our model simulates
a strengthening of the South American Summer Monsoon over
Brazil, in agreement with the d18O record from Botuverá cave (Cruz
et al., 2005; Wang et al., 2007) (Fig. 9c).
5. Discussion and conclusion
The transient deglaciation experiments performed in this study
successfully capture the dominant climate and some biochemical
features of the last deglaciation in both hemispheres. This good
correspondence between model results and paleo-proxy data
suggests that most of the millennial-scale variability recorded during
the last deglaciation can be explained by freshwater-induced variations in the deep water formation strength in the North Atlantic,
Southern Ocean and North Pacific. Based on our modeling results it is
justified to conclude that the North Atlantic is the key driver of
millennial-scale variability during the Last Glacial Termination.
In this study freshwater was mostly added in a broad area of the
North Atlantic, not taking into account differences that would arise
from changes in the freshwater routing around the North Atlantic
(e.g. through the Norwegian Sea, Labrador Sea, Gulf of Mexico)
(Aharon, 2003; Tarasov and Peltier, 2006; Sionneau et al., 2010).
Therefore, we did not differentiate between the North Atlantic deep
water formation sites, i.e. the Labrador Sea and the Greenland Iceland
and Norwegian (GIN) Sea. As seen in Fig.1a, paleo-proxy records from
different North Atlantic locations suggest slightly different scenarios.
The Bermuda rise 231Pa/230Th data (McManus et al., 2004) and the
South Iceland rise ventilation age record (Thornalley et al., 2011)
indicate a weakened AMOC already at 18 ka B.P., whereas the Iberian
margin 231Pa/230Th data shows a weakened AMOC starting 16 ka B.P
(Gherardi et al., 2005). Gherardi et al. (2005) suggested that the
Labrador Sea NADW formation site might have weakened before the
GIN site, which is actually not supported by the data presented in
Thornalley et al. (2011). In the Northern Hemisphere, the Cariaco
basin reflectance record (Peterson et al., 2000) is the only dataset
indicating a late (w16.2 ka B.P.) initiation of H1. On the other hand,
the South Iceland rise data suggest a greater ventilation between 16
and 15 ka B.P., which is not seen in other North Atlantic circulation
and temperature records. Even if a strong weakening of the NADW
formation most likely occurred during H1, the impact of variations in
the North Atlantic convection sites needs to be further investigated.
Whether a more detailed knowledge of the freshwater routing into
the North Atlantic would help to reconcile the conflicting timing in
paleo-proxy data during H1 remains to be shown.
A shut down of the AMOC during H1 leads to cold conditions
over the North Atlantic while the Southern Hemisphere starts to
warm. A related reduction of Atlantic-Pacific moisture transport
leads to an increase of North Pacific salinity and eventually triggers
deep water Formation in the Bering Sea. This leads to the establishment of a deep Pacific Meridional Overturning Circulation
during H1 in agreement with paleo-proxy records (Okazaki et al.,
2010). The closing of the Bering Strait is necessary for the formation of NPDW (Okumura et al., 2009). Further North Pacific paleoSST records and modeling studies would be needed to better
constrain the associated SST patterns in the North Pacific during H1.
The rapid and strong resumption of the AMOC at about 15 ka B.P.
leads to the Bølling warming in the North Atlantic and initiates the
ACR at high Southern latitudes. The ACR can be simulated by
a weakening of the AABW formation, either as a bipolar seesaw
response to the strengthened AMOC or as a result of a freshwater
input in the Southern Ocean. The knowledge of AABW dynamics is
too limited at the moment to conclude whether the bipolar seesaw
is sufficient to weaken the AABW or to estimate the amount of
freshwater necessary to weaken the AABW for about 2000 years. In
agreement with paleoproxies (Sachs and Anderson, 2005;
Anderson et al., 2009), our model successfully simulates
a decrease in export and opal production in the Southern Ocean
during the ACR due to the sea-ice edge advance and colder conditions. It should be however noted that our model does not include
iron limitation. Changes in export production therefore do not take
into account any possible changes in iron supply to the euphotic
zone as well as non-Redfield processes.
Concurrently in the North Hemisphere, the Older Dryas
(14.4e14 ka B.P) is simulated by a substantial weakening of the
AMOC. Our simulation thus suggests that melting from both,
Antarctica and the Northern Hemispheric ice sheets contributed to
MWP-1A at w14.4 ka B.P. The exact timing, shape and location of
the different freshwater pulses contributing to MWP-1A still need
to be better constrained to allow for a more quantitative comparisons between climate model simulations and paleo-proxy data.
The model results suggest that climate sensitivity to southern
ocean freshwater input is largest in the high southern latitudes.
Future model-proxy comparisons could thus improve the estimates
of the Antarctic meltwater contribution to MWP-1A.
The Younger Dryas is represented by a freshwater input into the
Arctic Ocean following the hypothesis of Tarasov and Peltier (2005)
and Murton et al. (2010). The induced shut down of the AMOC
captures most but not all of the features of the YD in the Northern
Hemisphere. 231Pa/230Th data of (McManus et al., 2004) suggests
that only a weakening and not a complete shut down of the AMOC
occurred during the YD. Our model is not able to simulate an
intermediate strength of the AMOC. In addition, the amplitude of
the cooling observed in GISP2 (Alley, 2000), in the North Atlantic
(Cacho et al., 1999, 2001; Bard et al., 2000; Calvo et al., 2001;
Waelbroeck et al., 2001; Bard, 2002; Martrat et al., 2004, 2007)
and in the Cariaco basin (Lea et al., 2003) can only be obtained in
our model as the consequence of a shut down of the AMOC. In the
Southern Hemisphere, temperatures increase in phase with the
strengthening of AABW formation, starting around 13 ka B.P. We
therefore find that, similarly to H1, temperature and precipitation
variations between the North Atlantic and the Southern Hemisphere are anti-correlated during the YD, in conformance with the
bipolar seesaw mechanism.
Fig. 12 shows the time series of simulated changes in terrestrial
primary production in different latitudinal bands compared to the
atmospheric CH4 as recorded in GISP2 ice core (Brook et al., 2000).
Wetlands are a main source of atmospheric CH4 and it has been
estimated that tropical region wetlands (20 Ne30 S) were
responsible for about 60% of the total wetland emissions, while the
high northern latitudes wetlands (45 Ne70 N) were responsible
for about 38% (Bartlett and Harris, 1993). Due to the cold and dry
conditions prevailing in most of the Northern Hemisphere during
H1, terrestrial primary production slightly decreases in the latitudinal band 0e20 N (10%). Wetter and warmer conditions in
the Southern Hemisphere lead to greater terrestrial primary
production in the latitudinal band 0e30 S (þ20%), which could
induce the measured atmospheric CH4 increase. Overall, the
simulated terrestrial carbon stock during H1 increases by about 100
GtC. The abrupt increase in atmospheric CH4 at 15 ka B.P. corresponds to the Bølling warming and the associated warmer and
wetter conditions in the Northern Hemisphere. The terrestrial
primary productivity increases by w20% in the latitudinal band
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L. Menviel et al. / Quaternary Science Reviews 30 (2011) 1155e1172
0e20 N and 35% over 45 Ne70 N (Fig. 12). As a result of the drier
conditions prevailing in the Southern Hemisphere during the ACR,
the terrestrial primary production slightly decreases (5%) in the
latitudinal band 0e30 S. Overall, the terrestrial carbon stock
increases by about 260 GtC between 15 and 13 ka B.P. Finally, the
drier conditions over the Northern Hemisphere during the YD
induce a decrease in terrestrial primary production by 60% in the
latitudinal band 0e20 N and by 30% in the latitudinal band
45e70 N. The terrestrial carbon stock decreases by only 60GtC
during the YD after which it increases by almost 300GtC.
In this study we suggest a sequence of changes in the strength of
deep water formation in the North Atlantic, the Southern Ocean and
the North Pacific, which gives a reasonable representation of the
reconstructed millennial-scale climatic variability during the Last
Glacial Termination. To a large extent, we find that the relative
timing of events observed in paleo-proxy data from different
regions is consistent with the relative timing in the model.
Ensemble-simulations with some form of data-assimilation (Goosse
et al., 2010b; Jackson et al., 2010) could help to put an improved
constraint on the freshwater forcing time series and, subsequently,
the phase relations among proxy time series. A more detailed study
of the climatic influence of the different deep water formation sites
in the North Atlantic (similar to (Roche et al., 2010)) would also help
to improve the understanding of the deglacial climatic changes.
Acknowledgment
The authors would like to acknowledge the NOAA NGDC database and thank O. Heiri, K. Holmgren, K. Pahnke and J. Sachs for
providing their datasets. The authors would like to thank P. Baker for
discussion about his dataset. We thank W. Hiscock for his graphic
support. This research was supported by NSF grant No. 1010869.
Additional support was provided by the Japan Agency for MarineEarth Science and Technology (JAMSTEC), by NASA and NOAA
through their sponsorship of research activities at the International
Pacific Research Center. A. Mouchet acknowledges support from the
Belgian Science Policy (BELSPO contract SD/CS/01A). This is IPRC
publication number 756 and SOEST contribution number 8096.
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