Lam2007a
Earth and Planetary Science Letters 259 (2007) 400 – 413
www.elsevier.com/locate/epsl
Modulation of the bipolar seesaw in the Southeast Pacific
during Termination 1
Frank Lamy a,⁎, Jérôme Kaiser b , Helge W. Arz b , Dierk Hebbeln c , Ulysses Ninnemann d ,
Oliver Timm e , Axel Timmermann e , J.R. Toggweiler f
a
f
Alfred-Wegener-Institute for Polar and Marine Research, Am Alten Hafen 26, 27568 Bremerhaven, Germany
b
GeoForschungsZentrum-Potsdam, Telegrafenberg, 14473 Potsdam, Germany
c
MARUM – Center for Marine Environmental Sciences, University of Bremen, Leobener Strasse, 28359 Bremen, Germany
d
Bjerknes Centre for Climate Research, University of Bergen, Allégaten 55, 5007 Bergen, Norway
e
IPRC, SOEST, University of Hawaiʻi at Manoa, 2525 Correa Road, Honolulu, HI 96822, USA
Geophysical Fluid Dynamics Laboratory, National Oceanic and Atmospheric Administration, P.O. Box 308, Princeton, NJ 08542, USA
Received 30 January 2007; received in revised form 25 April 2007; accepted 26 April 2007
Available online 8 May 2007
Editor: H. Elderfield
Abstract
The termination of the last ice age (Termination 1; T1) is crucial for our understanding of global climate change and for the
validation of climate models. There are still a number of open questions regarding for example the exact timing and the
mechanisms involved in the initiation of deglaciation and the subsequent interhemispheric pattern of the warming. Our study is
based on a well-dated and high-resolution alkenone-based sea surface temperature (SST) record from the SE-Pacific off southern
Chile (Ocean Drilling Project Site 1233) showing that deglacial warming at the northern margin of the Antarctic Circumpolar
Current system (ACC) began shortly after 19,000 years BP (19 kyr BP). The timing is largely consistent with Antarctic ice-core
records but the initial warming in the SE-Pacific is more abrupt suggesting a direct and immediate response to the slowdown of the
Atlantic thermohaline circulation through the bipolar seesaw mechanism. This response requires a rapid transfer of the Atlantic
signal to the SE-Pacific without involving the thermal inertia of the Southern Ocean that may contribute to the substantially more
gradual deglacial temperature rise seen in Antarctic ice-cores. A very plausible mechanism for this rapid transfer is a seesawinduced change of the coupled ocean–atmosphere system of the ACC and the southern westerly wind belt. In addition, modelling
results suggest that insolation changes and the deglacial CO2 rise induced a substantial SST increase at our site location but with a
gradual warming structure. The similarity of the two-step rise in our proxy SSTs and CO2 over T1 strongly demands for a forcing
mechanism influencing both, temperature and CO2. As SSTs at our coring site are particularly sensitive to latitudinal shifts of
the ACC/southern westerly wind belt system, we conclude that such latitudinal shifts may substantially affect the upwelling of
deepwater masses in the Southern Ocean and thus the release of CO2 to the atmosphere as suggested by the conceptual model of
[Toggweiler, J.R., Rusell, J.L., Carson, S.R., 2006. Midlatitude westerlies, atmospheric CO2, and climate change during ice ages.
Paleoceanography 21. doi:10.1029/2005PA001154].
© 2007 Elsevier B.V. All rights reserved.
Keywords: paleooceanography; Termination 1; Southeast Pacific; bipolar seesaw; alkenones
⁎ Corresponding author. Tel.: +4947148312121; fax: +4947148311923.
E-mail address: [email protected] (F. Lamy).
0012-821X/$ - see front matter © 2007 Elsevier B.V. All rights reserved.
doi:10.1016/j.epsl.2007.04.040
F. Lamy et al. / Earth and Planetary Science Letters 259 (2007) 400–413
1. Introduction
The termination of the last ice age (Termination 1;
T1) is the last major climate transition of the Earth's
recent geological history and is thus crucial for our
understanding of recent climate processes and the
validation of climate models. Though T1 is accordingly
very well studied involving numerous proxy records
from both marine and terrestrial archives (e.g., Alley and
Clark, 1999; Clark et al., 1999, 2004; Rinterknecht
et al., 2006) as well as modelling studies (e.g., Knorr
and Lohmann, 2003; Weaver et al., 2003), there are still
a number of open questions regarding for example the
exact timing and the mechanisms involved in the
initiation of deglaciation and the subsequent interhemispheric pattern of the warming. Based on the Milankovitch concept, the ultimate drivers for the glacial
termination are the increase in Northern Hemisphere
(NH) summer insolation and non-linear responses from
continental ice-sheets and particularly atmospheric
greenhouse gases such as CO2 that transfer the northern
signal globally (e.g. Clark et al., 1999). However, it has
also been repeatedly suggested that the Southern
Hemisphere (SH) leads the deglaciation and warming
in the NH (e.g., Bard et al., 1997), whereas a reevaluation of available ice-core and marine records
covering T1 (Alley et al., 2002) suggests a northern
temperature lead on orbital time-scales.
Part of the divergent views on possible interhemispheric leads or lags during T1 arise from the
pronounced millennial-scale variations that are superimposed on primarily insolation-driven orbital-scale
changes and are markedly different between the NH and
SH. The general warming trend that may start as early as
23,000 years before present (23 kyr BP), based on
Greenland and Antarctic ice-core records (e.g., Alley
and Clark, 1999; Blunier and Brook, 2001), is further
accentuated between ∼ 17 and 19 kyr BP in the south,
whereas NH records show a return to cold conditions
that culminate at the time of Heinrich event (HE) 1 (e.g.,
Alley and Clark, 1999). Thereafter, NH temperature
abruptly increased into the Bølling/Allerød (B/A) warm
period. In Antarctica, the deglacial warming trend was
partly interrupted by a millennial-scale cooling event
(Antarctic Cold Reversal, ACR) that began around the
time of the B/A warming and ended close to the
beginning of the Younger Dryas (YD) cold phase
observed in the NH (e.g., Blunier and Brook, 2001;
Morgan et al., 2002). The present picture of climate
pattern during T1 is thus largely focussed on high
latitude records in particular from Greenland and
Antarctic ice-cores that have been synchronized by
401
correlating globally recordable methane fluctuations
(e.g., Blunier and Brook, 2001; Morgan et al., 2002;
Epica Community Members, 2006). However, this
correlation reveals ambiguities over the interval of the
beginning deglacial warming in the SH making the
analysis of interhemispheric climate pattern over this
important interval more difficult.
Marine records from the SH have been involved to a
much lesser extent. The available data from the
Southern Ocean (e.g., Bianchi and Gersonde, 2004;
Shemesh et al., 2002) and southern mid-latitudes (e.g.,
Pahnke et al., 2003) are generally consistent with the
Antarctic records but dating uncertainties are high due to
scarce datable material and/or large and potentially
variable 14C reservoir ages. In addition, an increasing
number of high-resolution records from the tropics have
recently become available (e.g., Lea et al., 2006; Visser
et al., 2003). As deglacial warming in some of these
records occurred largely in phase with the CO2 increase
as observed in Antarctic ice-cores, they have been
interpreted in support for a tropical “trigger” for the
deglaciation (Visser et al., 2003).
In this paper, we attempt to better understand the
sequence of events over the last termination, on an
absolute time-scale, based on a new sea surface
temperature (SST) record from the SE-Pacific with
exceptional time-resolution and dating accuracy over T1
(i.e., 10–25 kyr BP). Our SST data are from Ocean
Drilling Project (ODP) Site 1233 located at the southern
Chilean continental margin at 41°S within the northernmost reach of the Antarctic Circumpolar Current (ACC)
and the southern westerly wind belt (Fig. 1). In previous
works, we showed that the complete ∼70-kyr-long
alkenone SST record at Site 1233 closely follows
millennial-scale temperature fluctuations as observed
in Antarctic ice cores (Kaiser et al., 2005; Lamy et al.,
2004). However, the absolute age-scale over the earlier
part of the last glacial, when large amplitude methane
fluctuations allow a detailed inter-correlation of Greenland and Antarctic ice-cores (Blunier and Brook, 2001;
Epica Community Members, 2006), is less well defined
in marine sediments due to increasing uncertainties in
radiocarbon dating and calendar year conversion. Therefore, we now substantially increased the time resolution
around T1, an interval that spans ∼27 m composite core
depth at Site 1233, and added a number of new 14C AMS
dates, now with an average spacing of ∼ 1200 years.
2. Investigation area
Site 1233 (41°00′S; 74°27′W) is located 38 km
offshore (20 km off the continental shelf) at 838 m water
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F. Lamy et al. / Earth and Planetary Science Letters 259 (2007) 400–413
depth in a small forearc basin on the upper continental
slope off Southern Chile (Fig. 1) away from the pathway
of major turbidity currents (Mix et al., 2003). The region
is located within the northernmost reach of the Antarctic
Circumpolar Current (ACC) at the origin of the Peru–
Chile Current (PCC) (Fig. 1). The ACC brings cold,
relatively fresh, nutrient-rich, Subantarctic Surface
Water originating from the region north of the
Subantarctic Front. The northern part of the ACC splits
around ∼43°S into the PCC flowing northward and the
Cape–Horn Current (CHC) turning towards the south
(Strub et al., 1998). The mean annual SST at ∼ 41°S
(ODP Site 1233) is ∼ 14 °C and varies between ∼ 11 °C
in winter and ∼ 16 °C in summer, i.e. with a seasonal
amplitude of ∼5 °C. Linked to the northern boundary of
the ACC, steep latitudinal SST gradients occur south of
Site 1233, a region that is increasingly influenced by the
southern westerly wind belt (Fig. 1). Northward, the
SST isotherms take a more meridional orientation,
primarily as a result of the equatorward advection of
cold water in the PCC and to a lesser extent as a direct
consequence of increasing coastal upwelling towards
the central and northern Chilean margin (Tomczak and
Godfrey, 2003).
3. Material, methods, and chronology
3.1. Sampling
Five Advanced Piston Corer holes were drilled at Site
1233 to ensure a complete stratigraphic overlap between
cores from different holes. Detailed comparisons
between high-resolution core logging data performed
shipboard demonstrated that the complete sedimentary
sequence down to 116.4 meters below surface (mbsf)
was recovered. Based on these data, a composite
sequence (the so-called splice) was constructed representing 135.65 meters composite depth (mcd). Discrete
samples for alkenone analyses were taken from the
interval that covers Termination 1 (T1) (see age model)
with an average resolution of ∼ 15 cm resulting in a
temporal resolution of ∼ 90 years of our alkenone SST
record. Additional samples for 14C accelerator mass
spectrometry (AMS) dating were taken from the splice
and, in some cases, from outside the splice.
3.2. Age model
In this study we present data from the composite
sequence between 10 and 5 thousand calendar years
Fig. 1. (A) Map of the South Pacific Ocean and adjacent areas showing
major surface currents (after Tomczak and Godfrey, 2003) and the
location of marine sediment cores (TR163-19 from the tropical eastern
Pacific (Lea et al., 2006; Spero and Lea, 2002); ODP 1233 from the
Southeast Pacific (this study)) and ice-cores (Byrd (Blunier and Brook,
2001); Dome C (Epica Community Members, 2004); Dronning Maud
Land (DML) (Epica Community Members, 2006); Law Dome
(Morgan et al., 2002); Taylor Dome (Indermühle et al., 2000))
discussed in the paper. (B) Annual mean SST (°C) distribution in the
Southeast Pacific between 30°S and 60°S (NOAA–CIRES Climate
Diagnostics Center (http://www.cdc.noaa.gov/index.html)) and location of ODP Site 1233 at the northern margin of the ACC. The contour
interval is 0.5 °C. Further shown is a simplified view of the major
current systems (PCC = Peru–Chile Current; ACC = Antarctic Circumpolar Current; CHC = Cape Horn Current).
F. Lamy et al. / Earth and Planetary Science Letters 259 (2007) 400–413
before present (kyr BP) representing ∼ 13 mcd to
∼ 40 mcd. The age model of this interval is based on
thirteen 14C AMS datings (Table 1) with an average
spacing of ∼ 1200 years and linear interpolation between
the dates. 14C ages were primarily calibrated with the
INTCAL04 calibration curve (Reimer et al., 2004).
However, the INTCAL04 calibration curve is poorly
constrained for the interval between ∼ 12,500 and
∼ 14,500 14C yr BP with few data points and missing
surface coral data (Robinson et al., 2005) (Fig. 2).
Radiocarbon data from the Cariaco basin (Hughen et al.,
2004) suggest the presence of a radiocarbon plateau
lasting from ∼ 12,900 to ∼13,300 14C yr BP (∼ 15.7 to
∼ 17 kyr BP) (Fig. 2). Similar results have been obtained
from a densely 14C-dated marine sediment core in the
Northwest Pacific (Sarnthein et al., 2006). Likewise, two
of our 14C AMS datings (at 21.39 mcd and 23.69 mcd)
within this interval revealed 14C-ages very close together
that would result in anomalously high sedimentation
rates (Table 1). Therefore, we applied here the CalPal_SFCP_2005 (www.calpal.de) calibration curve
which is primarily based on the Cariaco basin record
(Hughen et al., 2004) in this interval and contains a
number of data points (Fig. 2). For the CalPal_SFCP_2005 calibration curve, the original GISP2synchronized gray-scale record has been adapted to the
Greenland time-scale of Shackleton et al. (2004). This
has been done by linearly interpolating between the
unchanged base of the Bølling/Allerød at 14.66 kyr BP
and the base of Greenland interstadial at 29 kyr BP
(compared to 27.84 kyr BP in the original synchronization to GISP2 (Hughen et al., 2004)). Within the for this
study relevant interval the offset is however very small.
Table 1 compares the calibrated ages using different
calibration curves (including CalPal_SFCP_2005,
INTCAL04, and the most recent Fairbanks U/Th-based
calibration curve (Fairbanks et al., 2005)). Except for
the above mentioned interval, the calibrated ages are
nearly indistinguishable making the timing of the first
and second major warming step discussed in this paper
very robust. Calendar ages derived with the CalPal_SFCP_2005 calibration curves are however significantly older for three datings within the poorly
constrained interval in the coral-based calibration
curves. The presence of the above mentioned radiocarbon plateau induces a significant uncertainty in the
calibrated ages of the dating at 23.69 mcd. However, the
sedimentation-rates achieved by calibrating the 14Cdatings with the CalPal_SFCP_2005 calibration curve
appear more realistic assuming only moderately variable
sedimentation rates at Site 1233 as shown by the other
datings (Table 1).
403
As discussed in detail in our previous publications
(Kaiser et al., 2005; Lamy et al., 2004), we assume no
regional deviation from the global reservoir effect of
∼ 400 years because of the presence of an early
Holocene volcanic ash layer at Site 1233 (that has
been likewise dated on land) and the position of our site
significantly south of the Chilean upwelling zone (Strub
et al., 1998) and north of the southern polar front where
higher reservoir ages may be expected. In addition, the
new datings presented in this paper that fall on the above
mentioned radiocarbon plateau support the ∼ 400 years
reservoir age assumption. Larger reservoir ages would
move the corrected 14C-ages after the plateau (Fig. 2)
and would thus yield anomalous sedimentation rates
(Table 1).
We revised all radiocarbon-based age models of
published studies shown in our paper as outlined above
for our own datings. All ice-core age models are plotted
on the new Greenland Ice Core Chronology 2005
(GICC05) that is based on annual layer counting back to
42 kyr BP (Andersen et al., 2006). The Epica Dronning
Maud Land (DML) and Byrd ice-cores have been
synchronized to the Greenland record using the pattern
of millennial-scale methane fluctuations (Epica Community Members, 2006; Blunier et al., 2007). The iceage scale of the Epica Dome C (Dome C) record has
been synchronized to that of DML using volcanic and
dust tie points based on continuous sulfate, electrolytic
conductivity, dielectric profiling, particulate dust, and
Ca2+ data available for both cores (Epica Community
Members, 2006). A gas-age model based on the
GICC05 for Dome C has not yet been published. We
therefore synchronized the Dome C and DML methane
records taking a minimum number of tie-points at the
large fluctuations around the YD and B/A and a minor
peak in methane close to 23 kyr BP (Fig. 3).
3.3. Alkenone measurements
Alkenones were extracted from 1 to 3 g of freezedried and homogenized sediment following a procedure
described in detail by Müller et al. (1998). The extracts
were analysed by capillary gas chromatography using an
HP 5890 serie II Plus gas chromatograph equipped with
a 60-m column (J&W DB5MS, 0.32 mm × 0.1 μm),
split/splitless and flame ionization detection. Helium
was used as carrier gas with a constant pressure of
150 kPa. The oven temperature was programmed to
reach 50–250 °C at 25 °C/min, 250–290 at 1 °C/min,
followed by a plateau of 26 min, and 290–310 at 30 °C/
min, with the final temperature being maintained for
10 min.
404
Shown are the calendar ages and resulting linear sedimentation rates based on the CalPal_SFCP_2005 (www.calpal.de) primarily based on the GISP2-synchronized Cariaco basin record (Hughen
et al., 2004) INTCAL04 (Reimer et al., 2004), and the most recent Fairbanks U/Th-based calibration curve (Fairbanks et al., 2005). Marked in red is one dating that is affected by the radiocarbon
plateau shown by the CalPal_SFCP_2005 calibration curve (see Fig. 2 and discussion in Section 3.2).
⁎ Mark calibrated ages used for the final age model.
F. Lamy et al. / Earth and Planetary Science Letters 259 (2007) 400–413
Table 1
C ages obtained by accelerator mass spectrometry dating of mixed planktonic foraminifera samples (primarily Globigerinoides bulloides and Neogloboquadrina pachyderma), performed at the
Leibniz–Labor AMS facility in Kiel, Germany
14
F. Lamy et al. / Earth and Planetary Science Letters 259 (2007) 400–413
405
Fig. 2. Calibration of 14C AMS dates over Termination 1. 14C ages were primarily calibrated with the INTCAL04 calibration curve (Reimer et al.,
2004) (calibration results are indicated by black bars with probability distribution). However, the INTCAL04 calibration curve is poorly constrained
for the interval between ∼ 12,500 and ∼14,500 14C yr BP with few data points and missing surface coral data (Robinson et al., 2005). For this
interval, we used the CalPal_SFCP_2005 (www.calpal.de) calibration curve (green bars with probability distribution). Doted lines mark one date
that is substantially influenced by the radiocarbon plateau evident in the CalPal_SFCP_2005 that derives from the Cariaco basin data points
(Hughen et al., 2004) in this interval. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of
this article.)
Quantification of the alkenones was achieved using
2-nonadecanone (C19H38O) as internal standard and
HPGC ChemStation as analytical software. The alkenone unsaturation index UK′37 was calculated from
UK′37 = (C37:2) / (C37:3 + C37:2), where C37:2 and C37:3 are
the di- and tri-unsaturated C37 methyl alkenones. The
analytical precision was estimated to be ± 0.3 °C. For
conversion into temperature values, we used the culture
calibration of Prahl et al. (1988) (UK′37 = 0.034T +
0.039), which has been validated by core-top compilations (e.g., Müller et al., 1998). We assume that
alkenone-derived SST estimates at Site 1233 reflect
annual mean sea surface temperatures as suggested by
measurements on surface sediments at Site 1233 and
further north along the Chilean continental margin
(Kaiser et al., 2005; Kim et al., 2002). This does,
however, not exclude that alkenone SSTs could be
biased towards the spring bloom in productivity.
It has been recently observed that alkenones may be
substantially older than co-occurring planktic foraminifera (Mollenhauer et al., 2005). Holocene age differences
measured on the Site 1233 survey core GeoB 3313-1
showed rather constant age offsets of ∼ 1000 years
(Mollenhauer et al., 2005). Mollenhauer et al. (2005)
explained this offset as most likely resulting from
continuous resuspension/redeposition cycles induced
by internal tides and sediment focusing in morphologic
depressions such as the small basin at Site 1233. By
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F. Lamy et al. / Earth and Planetary Science Letters 259 (2007) 400–413
containing fine fraction and the coarse fraction foraminifera that have been used for dating in our study.
3.4. Modelling
Fig. 3. Illustration of the methane synchronization for Antarctic ice
core records. The Byrd, Dome C, and Dronning Maud Land (DML)
records are synchronized to the Greenland GRIP record (Epica
Community Members, 2006; Blunier et al., 2007) based on the layer
counted GICC05 (Andersen et al., 2006). Note that the methane
synchronization is uncertain for the interval from ∼ 17 to 23 kyr BP.
Triangles show tie points used for the methane synchronization of
DML and Dome C.
comparing the age offsets in different continental
margin settings, they further noted that age offsets
were largest where TOC contents and alkenone
concentrations are highest. Therefore, we expect that
the age offsets if they are indeed induced by resuspension/redeposition cycles should be much smaller for the
deglacial section where both TOC and alkenone
concentrations are significantly lower than during the
Holocene. Alkenone concentrations are in the order of
2000 to 3000 ng/g dry sediment during the Holocene
and 500 to 1000 ng/g dry sediment during the late
glacial. TOC contents range from Holocene values
between ∼1.5 and ∼ 2.5 wt.% to late glacial values
between ∼ 0.5 and ∼ 1 wt.% (Kaiser et al., 2005;
Martinez et al., 2006). We also note that Holocene grainsize data on the survey core GeoB 3313-1 suggest
constant and rather undisturbed fine-grained hemipelagic sedimentation (Lamy et al., 2001). Available
oceanographic data show that bottom water circulation
at the depth of Site 1233 (Antarctic Intermediate Water;
e.g., (Shaffer et al., 2004)) is rather too sluggish for the
re-suspension of sediments and internal waves have not
been described at the Chilean margin. We suggest that
the constant admixture of older material that would
affect the 14C ages of the alkenone fraction but not
significantly the reconstructed alkenone temperatures
would be likewise conceivable, a possibility that
Mollenhauer et al. (2005) did not exclude either. We
therefore assume that our SST record is not substantially
affected by any age offsets between the alkenone
The transient glacial–interglacial simulation shown
in Fig. 5H was performed with the ECBilt–Clio climate
model and includes the time-varying orographic and
albedo ice-sheet effects, greenhouse gas changes, and
orbital forcing variations. The time-varying greenhouse
gas forcing uses CO2, CH4, and N2O concentrations.
Concentration values were measured on the Antarctica
ice-core Taylor Dome (Indermühle et al., 2000; Smith
et al., 1999). The time-scale was aligned to the GISP2
time-scale (Meese et al., 1997). CH4 and N2O were
measured in samples from the GISP2 ice-core (Sowers
et al., 2003). The most relevant greenhouse gas changes
are associated with CO2. During the LGM the
estimated global radiative forcing anomaly with respect
to pre-industrial conditions amounts to about − 2 W/m2
(compared to − 0.22 W/m2 and − 0.25 W/m2 for CH4
and N2O, respectively).
4. Results and discussion
4.1. Sea surface temperatures off Chile compared to
Antarctic ice-core records
Deglacial warming in our alkenone SST record starts
at ∼ 18.8 kyr BP with a ∼ 2-kyr-long increase of nearly
5 °C until ∼ 16.7 kyr BP (Fig. 4A). Thereafter,
temperatures remain comparatively stable until the
beginning of a second warming step of ∼ 2 °C between
∼ 12.7 and ∼ 12.1 kyr BP. A comparison of our SST
record to different Antarctic ice-core records suggests a
general correspondence in the major temperature trends,
particularly the two-step warming over T1. As in our
SST record, in the Pacific Sector of Antarctica, (Byrd,
Fig. 4B) (Blunier and Brook, 2001), deglacial warming
initiated shortly after 19 kyr BP. Both records show very
similar millennial-scale variations before T1 though
these changes partly reveal larger offsets in particular
between ∼ 19 and 23 kyr BP where the methane
synchronization is uncertain (Blunier and Brook, 2001;
Epica Community Members, 2006) (i.e., the SST
minimum close to 22.5 kyr BP may well correspond to
the temperature minimum at ∼21.5 kyr BP in the Byrd
record). The new record from Dronning Maud Land
(Epica Community Members, 2006) (DML; Atlantic
Sector; Fig. 4C), shows a ∼600-year delayed initiation
of deglacial warming and a millennial-scale warming
between ∼23.5 and 24.5 kyr BP (Antarctic Isotope
F. Lamy et al. / Earth and Planetary Science Letters 259 (2007) 400–413
Fig. 4. Southeast Pacific SST record compared to different Antarctic
temperature proxy records over T1. (A) Alkenone SST record from
Site 1233 with radiocarbon datings (this study). (B) Oxygen isotope
record from the Byrd ice-core (Blunier and Brook, 2001) (coastal site,
Pacific sector). (C) Oxygen isotope record from the Dronning Maud
Land (DML) ice-core (Epica Community Members, 2006) (coastal
site, Atlantic sector). (D) Deuterium record from the Dome C ice-core
(Epica Community Members, 2004) (continental site in eastern
Antarctica). All ice-core records are plotted on the GICC05 displaying
100-year averages.
Maximum 2) that occurs ∼ 500 years later than a
warming shown in the SST data. In the Dome C record
(Epica Community Members, 2004) (continental site,
Fig. 4D), the initial warming starts at about the same time
as in the DML record. In general, the deglacial warming
as documented in Antarctic ice-cores is substantially
more gradual than observed in our SST record where
most of the initial warming occurs over a time-interval of
only ∼ 1200 years (∼ 18.8 to 17.6 kyr BP).
4.2. The bipolar seesaw
Millennial-scale temperature changes in Antarctica
over the last glacial may be consistently explained by the
bipolar seesaw concept that suggests an out-of-phase
millennial-scale climate pattern between the NH and SH
during the last glacial (e.g., Stocker et al., 1998). The
concept was later extended by including a time constant
that describes the thermal storage effect of the Southern
Ocean and explains why glacial Antarctic and Greenland
407
temperatures are not strictly anti-correlated but are rather
characterised by a lead–lag relationship (Knutti et al.,
2004; Siddall et al., 2006; Stocker and Johnsen, 2003)).
Over T1, the bipolar seesaw concept and its interference
with orbital-scale changes in different components of the
climate system has been much less investigated (Clark
et al., 2002). The initiation of deglacial warming in the
SE-Pacific shortly after 19 kyr BP coincides very closely
with a starting slowdown of the Atlantic meridional
overturning circulation (AMOC) (McManus et al., 2004)
(Fig. 5B) and a beginning NH (Greenland) cooling
towards HE 1 (Fig. 5C). The freshwater input that started
the reduction of the AMOC likely originated from the
beginning deglaciation of the NH ice sheets. This is
shown for example by a recent study on changes in the
extension of the Scandinavian Ice Sheet over T1
(Rinterknecht et al., 2006) (Fig. 5A) suggesting that
deglaciation began at ∼ 19 kyr BP synchronous with a
previously suggested sea-level rise in the order of 10 to
15 m (Clark et al., 2004). The related freshwater input
was later reinforced during subsequent HE 1 (Clark et al.,
2004; Rinterknecht et al., 2006). By inducing Antarctic
ice-sheet melting, the SH warming may then have fed
back to the NH by resuming the AMOC leading into the
B/A warm period as indicated by modelling studies
(Knorr and Lohmann, 2003; Weaver et al., 2003). During
this time interval, SH warming slowed down or even
reversed (ACR). Both our record and the Antarctic icecore data reveal that the second reduction of the AMOC
again reinforced the SH warming as documented in the
second major warming step during the NH YD cold
phase (Fig. 5).
The timing of both the initial and the second warming
step in our data, suggests that the SST response in the
mid-latitude SE-Pacific occurred quasi instantaneous to
the starting slowdown of the AMOC (Fig. 5). The
conceptual model of Stocker and Johnsen (2003) shows
such strict antiphase behaviour for the South Atlantic.
However, the occurrence of an “immediate” and high
amplitude response in our SST record requires a rapid
transfer of the Atlantic signal to the SE-Pacific without
involving the thermal inertia of the Southern Ocean that
contributed to the substantially more gradual and partly
delayed (in case of the DML and Dome C records)
deglacial temperature rise seen in Antarctic ice-cores.
The most plausible mechanism for this rapid transfer is a
seesaw-induced change of the coupled ocean–atmosphere system of the ACC and the southern westerly
wind belt. Using a coupled atmosphere–ocean–sea ice
model, Timmermann et al. (2005) show a substantial
decrease of westerly airflow between ∼ 40 and ∼ 50°S
and an increase further south in the South Pacific for an
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F. Lamy et al. / Earth and Planetary Science Letters 259 (2007) 400–413
AMOC shutdown experiment compared to a Last
Glacial Maximum simulation (Fig. 6). This latitudinal
shift of the SH westerlies is a robust feature in a number
North Atlantic water-hosing experiments with Coupled
General Circulation Models (Timmermann et al.,
in press).
4.3. Other forcings beyond the bipolar seesaw
The SST response to a weakening of the AMOC in
these and other model simulations (e.g., Knutti et al.,
2004; Schmittner et al., 2002) is, however, much smaller
than the initial warming observed at Site 1233 (Fig. 6).
Part of the high amplitude SST response at our Pacific
site is likely caused by the very pronounced regional
SST gradients (Fig. 1). These gradients are intimately
linked to the northern margin of the westerlies and the
ACC and provide a regional sensitivity that may not be
captured by the comparatively coarse climate models.
Moreover, in contrast to the glacial period, bipolar
seesaw induced climate variations over T1 are more
strongly superimposed by changes in important forcing
factors such as insolation and atmospheric CO2 content.
A transient glacial–interglacial simulation with the
ECBilt–Clio climate model which neglects late glacial
millennial-scale meltwater forcing, suggests a substantial SST rise in the SE-Pacific that likewise starts at
∼ 19 kyr BP (Fig. 5H). In this transient simulation, the
SE-Pacific temperature response to orbital and greenhouse gas forcing shows a more gradual increase in
contrast to the distinct two-step warming observed in
our record. This further strengthens the importance of
the superposition of seesaw related processes with other
forcings such as orbitally induced seasonal variations of
incoming solar radiation and atmospheric CO2 that were
unique to T1.
An additional warming not considered in most
climate models may also be related to substantially
decreasing atmospheric dust contents as recorded in
Antarctic ice-cores. Dust contents in Antarctic ice
decrease notably to already Holocene levels during the
first major warming step recorded in our SST record
(Fig. 5F). A minor decrease observed in the log-scaled
record (Fig. 5F) also occurs over the second major
warming step. Antarctic dust primarily originates from
Southern Patagonia and its content in the ice is controlled
Fig. 5. Compilation of paleoclimatic records to explain interhemispheric climate pattern over T1. (A) Time–distance diagram of
fluctuations of the southern Scandinavian ice-sheet (SIS) margin
(Rinterknecht et al., 2006) including the position of the 19-kyr sealevel rise after Clark et al. (2004). (B) 231Pa/230Th record from a
subtropical North Atlantic sediment core with radiocarbon datings
taken as a proxy for the strength of the Atlantic meridional overturning
circulation (McManus et al., 2004). (C) Oxygen isotope record of the
Gisp2 ice-core, Greenland (Grootes et al., 1993). (D) Alkenone SST
record from Site 1233 with radiocarbon datings. (E) CO2 record from
the Dome C ice-core (Monnin et al., 2001) (methane-synchronized to
the GICC05). (F) Atmospheric dust content record from the Dome C
ice-core (Delmonte et al., 2002) (on the GICC05). Small insert figure
shows dust record on log-scale for the interval 10–14 kyr BP. (G)
Carbon isotope record from Site TR163-19, eastern equatorial Pacific
(Spero and Lea, 2002). (H) Modelled SST record at Site 1233
conducted with a transient glacial–interglacial simulation (ECBilt–
Clio) including orographic and albedo ice-sheet effects, CO2 changes,
and orbital forcing.
F. Lamy et al. / Earth and Planetary Science Letters 259 (2007) 400–413
409
Fig. 6. Atmospheric response to a transient glacial meltwater experiment leading to a complete shutdown of the AMOC performed with the ECBilt–
Clio climate model (Timmermann et al., 2005). Shown is the difference of time-averaged wind stress (vectors, eastward direction means decrease of
westerly airflow) and temperature (shading) fields between the meltwater experiment and a Last Glacial Maximum simulation. This situation
represents the response to the slowdown of the AMOC beginning at ∼ 19 kyr BP as seen in the proxy records (Fig. 4). Note the decrease of westerly
airflow in the Southern Hemisphere mid-latitudes and increase in the Southern Ocean consistent with a latitudinal shift of the westerly wind belt. The
temperature response is, however, only minor.
by atmospheric circulation pattern over Patagonia and
the Southern Ocean in addition to impacts of regional
aridity and sea level (Delmonte et al., 2002; Wolff et al.,
2006). Though the climatic impact of regional atmospheric dust content changes are discussed controversially (Harrison et al., 2001), a slight warming in the
order of 0.5 to 1 °C in the SH mid and high latitudes in
response to decreasing atmospheric dust content levels
over T1 as indicated by climate models (Schneider von
Deimling et al., 2006) can not be excluded.
4.4. Southeast Pacific SSTs and atmospheric CO2
A two-step pattern as in our SST record is also
apparent in the CO2 record from the Dome C ice-core
(Monnin et al., 2001) that, however, results in the model
simulation only in a gradual warming (Fig. 5H). The
correspondence of the deglacial pattern in SE-Pacific
SST and the CO2 record is remarkable. The initial
warming (∼5 °C) in our SST record slightly predates
(∼ 700 years) the most significant increase in CO2
(∼ 35 ppmv, interval I in Monnin et al. (2001)). The
second major warming step during the NH YD (∼ 2 °C)
in our SST record coincides with another CO2 increase
of ∼ 15 ppmv (first part of interval IV in Monnin et al.
(2001)) (Fig. 5D–E). Assuming that our record largely
reflects shifts of the coupled ACC/westerlies system,
this concurrence is consistent with the previously
suggested important role of such latitudinal shifts in
controlling atmospheric CO2 contents (Ninnemann and
Charles, 1997; Toggweiler et al., 2006). Based on a
general circulation model, Toggweiler et al. (2006)
showed that the equatorward shifted SH westerlies
during the glacial allowed more respired CO2 to
accumulate in the deep ocean. During glacial terminations, the southward moving westerlies reduced polar
stratification and enhanced upwelling of deepwater
masses around Antarctica that would then have released
large amounts of the stored CO2 to the atmosphere.
Such a mechanism is supported by the occurrence of a
pronounced δ13C minimum recorded in thermoclinedwelling foraminifera in the equatorial Pacific (Spero
and Lea, 2002). Within dating uncertainties, the onset
of this event during T1 coincides with the initiation of
SST warming and beginning southward movement of
the westerlies (Fig. 5G) and has been interpreted in
terms of a breakdown of surface water stratification and
renewed Circumpolar Deep Water upwelling in the
Southern Ocean (Spero and Lea, 2002). The ∼ 700-year
delayed beginning of the initial CO2 rise compared to
the SE-Pacific SST rise is probably related to the
uncertain methane synchronization during the beginning deglaciation (Fig. 3). This interpretation is
supported by the exact beginning of the second step
during the YD when the synchronization is very
accurate. In addition, new constraints on the gas age–
ice age difference along the Epica ice-cores suggest that
the lag of the CO2 increase at the start of T1 as
proposed by Monnin et al. (2001) is overestimated and
that the CO2 increase could well have been in phase or
slightly leading the temperature increase at Dome C
(Loulergue et al., 2007). This would move the initiation
of the CO2 rise close to the observed warming at our
site.
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F. Lamy et al. / Earth and Planetary Science Letters 259 (2007) 400–413
We observe a similar link between SE-Pacific SSTs
and CO2 for older intervals, for example the transition
from marine isotope stage (MIS) 4 to MIS 3, though in
this case a slightly lower CO2 increase of ∼ 25 ppmv
corresponds to a ∼ 5 °C SST increase (Fig. 7B–C). An
important question is why the partly substantial glacial
SST changes in the SE-Pacific and the associated shifts
of the SH westerlies and ACC system that resulted
sometimes in nearly similar CO2 changes as over the first
deglacial warming step have not initiated interglacial
conditions? One answer may be related to the duration of
the preceding cold phase. It is well conceivable that
larger amounts of CO2 were stored in the deep ocean
during the long-lasting glacial phase with low CO2
contents of late MIS 3 and MIS 2 compared to the
comparatively short MIS 4 that was preceded by nearly
interglacial conditions during late MIS 5. Thus, even
during comparable insolation changes, the release of
CO2 from the deep-water reservoir at T1 is expected to
have been larger. Probably more important is the particular combination of orbital-scale insolation changes
and millennial-scale climate variability over T1. NH
summer insolation similarly increased at the MIS 4/3
transition (Fig. 7A) and a major slowdown of the AMOC
(HE 6) likewise occurred during this interval (Fig. 7D).
However, the Toggweiler et al. (2006) model suggests
that the system may be characterised by a threshold
beyond that the westerlies and CO2 level would rapidly
move towards either their glacial or interglacial positions. It is well conceivable that this threshold was not
reached at the transition from MIS 4 to MIS 3 because
only one single major slowdown in the AMOC (i.e., HE
6) occurred late in the interval of increasing NH summer
insolation. Over T1, the beginning slowdown of the
AMOC towards HE 1 took place early in the interval of
insolation increase and was followed by a second slowdown (the YD) only a few millennia later (Fig. 7). Taken
together, both episodes likely moved the climate system
into interglacial conditions. The intervening resumption
of the AMOC during the B/A was apparently insufficient
to move the westerlies significantly back north as shown
by the longer SST “plateau” lasting from ∼ 16.7 to
∼ 12.7 kyr BP. Consistent with an interrupted rather than
reversed SH warming and southward movement of the
westerlies, the CO2 record of Dome C shows constant
values over the ACR (Fig. 5E). Furthermore, a clear
cooling during the ACR is likewise missing in the DML
ice-core record (Fig. 4C) but particularly well developed
in the more continental drilling sites as Dome C (Fig. 4D)
and Vostok (not shown on Fig. 4).
5. Conclusions
Fig. 7. Comparison of Southeast Pacific SST and atmospheric CO2
records over T1 to the transition from MIS 4 to MIS 3. (A) Summer
insolation at 60°N (Berger and Loutre, 1991). (B) CO2 record from the
Dome C (Monnin et al., 2001) (methane-synchronized to the GICC05)
and Taylor Dome ice-cores (Indermühle et al., 2000). The time-scale of
the Taylor Dome record has been adapted to the Gisp2-synchronized
age model of the Byrd ice-core (Blunier and Brook, 2001). (C)
Alkenone SST record from Site 1233 over the past 70 kyr (this study
and Kaiser et al., 2005) plotted on the time-scale as published in Kaiser
et al. (2005). (D) Oxygen isotope record of the Gisp2 ice-core,
Greenland (Grootes et al., 1993). Intervals of substantial SST warming
and CO2 increase over T1 and at the MIS4/3 transition are marked.
Numbers show approximate amplitude in ppmv and °C.
Our SE-Pacific SST record provides a unique opportunity to discuss globally relevant processes over Termination 1 on an absolute radiocarbon-based time-scale.
This point is particularly important as the lack of reliable
dating accuracy often hampered the exact dating of the
onset of deglacial warming in the Southern Ocean (due to
large and variable reservoir ages). Furthermore, Antarctic ice core records cannot be unambiguously synchronized to the Northern Hemisphere because of only minor
methane fluctuation during this particular interval.
Deglacial warming at the northern margin of the
Antarctic Circumpolar Current system (ACC) began
shortly after 19 kyr BP. Though this timing is largely
consistent with Antarctic ice-core records, the initial
warming in the SE-Pacific is more abrupt suggesting a
direct and immediate response to the slowdown of the
Atlantic thermohaline circulation through the bipolar
seesaw mechanism. This response requires a rapid tran-
F. Lamy et al. / Earth and Planetary Science Letters 259 (2007) 400–413
sfer of the Atlantic signal to the SE-Pacific without
involving the thermal inertia of the Southern Ocean that
may contribute to the substantially more gradual
deglacial temperature rise seen in Antarctic ice-cores.
The most plausible mechanism for this rapid transfer is a
seesaw induced change of the coupled ocean–atmosphere system of the ACC and the southern westerly
wind belt as supported by North Atlantic water-hosing
model experiments. The observed SST warming can
however not be explained by the bipolar seesaw alone.
Our modelling results suggest that a substantial part of
the signal is induced by insolation changes and the
deglacial CO2 rise that are superimposed on the bipolar
seesaw-induced signal but only lead to a gradual
warming at our site.
The similarity of the two-step rise in our proxy SSTs
and CO2 over T1 strongly demands for a forcing
mechanism influencing both, temperature and CO2. As
SSTs at our coring site are particularly sensitive to
latitudinal shifts of the ACC/southern westerly wind belt
system, we conclude that such latitudinal shifts may
substantially affect the upwelling of deepwater masses in
the Southern Ocean and thus the release of CO2 to the
atmosphere as suggested by the conceptual model of
Toggweiler et al. (2006). This connection of atmospheric
CO2 contents to SST changes in the Southeast Pacific
and the position of the westerlies may be very relevant
for our future climate as some models see significant
shifts of the westerlies under future greenhouse scenarios
(e.g., Yin, 2005).
An often discussed but still not resolved question is
the role of the tropics, in particular of the tropical Pacific.
Recent SST reconstructions from the Indo-Pacific Warm
Pool (Visser et al., 2003) and the eastern tropical Pacific
(Lea et al., 2006) show some similarities with our SST
record but with generally much smaller amplitudes. This
may well be explained by a transmission of South Pacific
SST warming through the surface ocean via the Eastern
Boundary Current system and through intermediate
water masses towards the tropics (Clark et al., 2004).
Such SST changes in the tropical Pacific may have
introduced important feedbacks by their large impact on
the hydrological cycle and the greenhouse gas concentration (Clark et al., 2004; Palmer and Pearson, 2003).
Acknowledgments
We thank P. Clark, A. Ganopolski, A. Mix, A.
Schmittner, B. Stenni, T. Stocker, J. Stoner, B. Weninger,
and R. Tiedemann for comments and suggestions as well
as T. Blunier, H. Fischer, and J. McManus for data. The
constructive reviews by M. Siddall and two anonymous
411
reviewers improved this paper. Financial support was
made available through the Deutsche Forschungsgemeinschaft (DFG). This research used samples provided
by the Ocean Drilling Program (ODP). The ODP is
sponsored by the U.S. National Science Foundation
(NSF) and participating countries under management of
Joint Oceanographic Institutions (JOI), Inc.
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