CPClimate of the PastCPClim. Past1814-9332Copernicus PublicationsGöttingen, Germany10.5194/cp-14-1405-2018Ice core evidence for decoupling between midlatitude atmospheric water
cycle and Greenland temperature during the last deglaciationIce core evidence for decouplingLandaisAmaëlleamaelle.landais@lsce.ipsl.frCapronEmilieMasson-DelmotteValériehttps://orcid.org/0000-0001-8296-381XToucanneSamuelRhodesRachaelhttps://orcid.org/0000-0001-7511-1969PoppTrevorVintherBoMinsterBénédictePriéFrédéricLaboratoire des Sciences du Climat et de l'Environnement, IPSL, UMR
8212, CEA-CNRS-UVSQ-UPS, Gif sur Yvette, FranceCentre for Ice and Climate, Niels Bohr Institute, University of
Copenhagen, Juliane Maries Vej 30, 2900, Copenhagen, DenmarkBritish Antarctic Survey, High Cross, Madingley Road, Cambridge, CB3
0ET, UKIFREMER, Laboratoire Géophysique et enregistrement
Sédimentaire, CS 10070, 29280 Plouzané, FranceDepartment of Earth Sciences, University of Cambridge, Downing
Street, Cambridge, CB2 3EQ, UKAmaëlle Landais (amaelle.landais@lsce.ipsl.fr)8October20181410140514153June201814June201817September2018This work is licensed under the Creative Commons Attribution 4.0 International License. To view a copy of this licence, visit https://creativecommons.org/licenses/by/4.0/This article is available from https://cp.copernicus.org/articles/14/1405/2018/cp-14-1405-2018.htmlThe full text article is available as a PDF file from https://cp.copernicus.org/articles/14/1405/2018/cp-14-1405-2018.pdf
The last deglaciation represents the most recent example of
natural global warming associated with large-scale climate changes. In
addition to the long-term global temperature increase, the last deglaciation
onset is punctuated by a sequence of abrupt changes in the Northern
Hemisphere. Such interplay between orbital- and millennial-scale variability
is widely documented in paleoclimatic records but the underlying mechanisms
are not fully understood. Limitations arise from the difficulty in
constraining the sequence of events between external forcing, high- and low-
latitude climate, and environmental changes.
Greenland ice cores provide sub-decadal-scale records across the last
deglaciation and contain fingerprints of climate variations occurring in
different regions of the Northern Hemisphere. Here, we combine new ice
d-excess and 17O-excess records, tracing changes in the
midlatitudes, with ice δ18O records of polar climate. Within
Heinrich Stadial 1, we demonstrate a decoupling between climatic conditions
in Greenland and those of the lower latitudes. While Greenland temperature
remains mostly stable from 17.5 to 14.7 ka, significant change in the midlatitudes of the northern Atlantic takes place at ∼16.2 ka, associated
with warmer and wetter conditions of Greenland moisture sources. We show that
this climate modification is coincident with abrupt changes in atmospheric
CO2 and CH4 concentrations recorded in an Antarctic ice
core. Our coherent ice core chronological framework and comparison with other
paleoclimate records suggests a mechanism involving two-step freshwater
fluxes in the North Atlantic associated with a southward shift of the
Intertropical Convergence Zone.
Introduction
The last deglaciation (∼19 000 to 11 000 years before
present, ka) is the most recent major reorganization of global climate and is
thus extensively documented by proxy records from natural climate archives.
The wealth of high-resolution records from well-dated archives and data
synthesis obtained over the past decades show two modes of climate
variability during this period (e.g., Denton et al., 2010; Clark et al.,
2012). The first is a long-term increase in global surface temperature and
atmospheric CO2 concentration between 18 and 11 ka. Superimposed
on this is a sequence of centennial-scale transitions between three
quasi-stable intervals documented in Northern Hemisphere temperature, namely
(i) the Heinrich Stadial 1 (∼17.5–14.7 ka) that encompasses the
massive rafting episode known as Heinrich event 1 (from ∼16 ka);
(ii) the Bølling–Allerød warming phase (∼14.7 to 12.9 ka); and
(iii) the Younger Dryas cold phase (∼12.9 to 11.7 ka). This three-step
sequence coincides with rapid variations in the Atlantic Meridional Oceanic
Circulation (AMOC) (McManus et al., 2004), with evidence for a weak
meridional overturning in the North Atlantic during the cold period
encompassing Heinrich Stadial 1 and the Younger Dryas.
Our understanding of the mechanisms at play during these North Atlantic cold
phases remains limited. First, recent studies challenge the earlier
attribution of the AMOC slowdown during Heinrich Stadial 1 to the impact of
the iceberg rafted debris (IRD) from the Laurentide ice sheet through Hudson
Strait (Álvarez-Solas et al., 2011). In particular, meltwater releases from
the European ice sheet occurring as early as 19 or 20 ka may have played an
important role in this AMOC slowdown (Toucanne et al., 2010; Stanford et al.,
2011; Hodell et al., 2017).
Second, major global reorganizations of the hydrological cycle have been
demonstrated during Heinrich Stadial 1. They can be separated into two phases.
In North America, a first time interval characterized by low lake levels
(referred to as the “big dry”, 17.5 to 16.1 ka) was followed by a second
time interval with high lake levels (referred to as the “big wet”, 16.1 to
14.7 ka) (Broecker et al., 2012), both apparently occurring during a stable
cold phase in Greenland temperature. The second phase of Heinrich Stadial 1
is also associated with a weak East Asian monsoon interval (Zhang et al.,
2014) understood to reflect a southward shift of the Intertropical
Convergence Zone (ITCZ). While there is growing evidence for large-scale
reorganizations of climate and low-latitude to midlatitude atmospheric water cycle
within Heinrich Stadial 1, the exact sequence of events is not known with
sufficient accuracy to understand the links between changes in North Atlantic
climate, AMOC, and the lower-latitude water cycle.
Linking changes in the high latitudes of the North Atlantic and the middle to
low latitudes requires precise absolute chronologies such as those obtained
from annual layer counting of Greenland ice (e.g., Andersen et al., 2006) or
U/Th dating of speleothems (e.g., Zhang et al., 2014). Unfortunately, absolute
dating uncertainties increase above 100 years during the last
deglaciation, precluding a direct comparison of proxy records at the
centennial scale. In this study, we circumvent this difficulty by using proxy
records measured on Greenland ice cores that represent both Greenland
temperature and midlatitude moisture source conditions.
Analytical method
Here, we present new water isotope records (δ18O,
d-excess =δD -8×δ18O,
17O-excess = ln(δ17O+1)-0.528×ln(δ18O+1)) from the NGRIP ice core (NGRIP community members, 2004)
reported on the annual-layer-counted Greenland Ice Core Chronology 2005
(hereafter GICC05, Rasmussen et al., 2006; Svensson et al., 2008) and
associated with relatively small absolute uncertainties over the last
deglaciation (maximum counting error of 100–200 years). Other Greenland and
Antarctic ice cores have been aligned on the GICC05 chronology, with a
maximum relative dating uncertainty of 400 years over the last deglaciation
(Rasmussen et al., 2008; Bazin et al., 2013; Veres et al., 2013).
Water stable isotope records (δ18O and d-excess, in
‰) from GRIP and NGRIP ice cores reported on the GICC05 chronology
(in thousands of years before year 2000 CE). (a) D-excess from
the NGRIP ice core (khaki: data obtained at INSTAAR SIL, Steffensen et al.,
2008; dark green: data obtained at LSCE, this study); d-excess from the
GRIP ice core (light green, Masson-Delmotte et al., 2005); (b)
d-excess from the NGRIP ice core after correction of the shift between the
INSTAAR SIL and LSCE (dark green) datasets, and d-excess from the GRIP ice
core (light green); C-δ18O from the NGRIP ice core (dark blue)
datasets, δ18O from the GRIP ice core (light blue). Grey
intervals display Heinrich Stadials (HS).
The new NGRIP δ18O and δD dataset was obtained at
Laboratoire des Sciences du Climat et de l'Environnement (LSCE) using a
Picarro laser cavity ring-down spectroscopy (CRDS) analyzer. The accuracy for
δ18O and δD measurements displayed here is about
0.1 ‰ and 1 ‰, respectively. This new dataset completes the
NGRIP high-resolution isotopic dataset published over the time period 11.5 to
14.7 ka with δ18O and δD measured respectively at the
University of Copenhagen and at the Institute of Arctic and Alpine Research
(INSTAAR) Stable Isotope Lab (SIL) (University of Colorado).
δ18O analyses were performed at the Niels Bohr Institute
(University of Copenhagen) using a CO2 equilibration technique
(Epstein et al., 1953) with an analytical precision of 0.07 ‰.
δD measurements at INSTAAR were made via an automated uranium
reduction system coupled to a VG SIRA II dual-inlet mass spectrometer (Vaughn
et al., 1998). Analytical precision for δD is ±0.5 ‰ or
better. Both time series show similar δ18O values, in agreement
with the reference δ18O series for NGRIP over the last
climatic cycle (NGRIP community members, 2004) within error bars. However,
while both LSCE and INSTAAR SIL d-excess time series display the same
3.5 ‰ decrease over the onset of Bølling–Allerød, the mean
d-excess level differs by 2.5 ‰ between the two records. Despite
several home standard intercalibrations between the two laboratories, this
difference remains unexplained and prevents any further discussion on the
absolute NGRIP d-excess levels. The new and published NGRIP d-excess
datasets are combined after a shift of the INSTAAR SIL d-excess series by
-2.5 ‰.
A synthesis of ice core records over the last deglaciation displayed
on the respective GICC05 and AICC2012 timescales with an identification of
two phases (1, orange box; 2, purple box) within Heinrich Stadial 1 (HS1)
as discussed in the text: we locate the transition between phases 1 and 2 at
the timing of the sharp increase in CO2 and CH4
concentrations, both being global atmospheric composition signals. The
Younger Dryas (YD) and Bølling–Allerød (BA) periods are also indicated.
(a) GRIP, NGRIP, and GISP2 δ18O (light blue, dark blue, and
black, respectively; Grootes et al., 1993; NGRIP community members, 2004)
interpolated at a 20-year resolution; (b) GRIP and NGRIP d-excess
(light and dark green, respectively; Jouzel et al., 2005, this study)
interpolated at a 20-year resolution; (c) NGRIP
17O-excess (orange curve shows the original series and the red
curve the 5-year running average; this study); (d) WAIS Divide
CH4 (Rhodes et al., 2015); (e) WAIS Divide CO2
(Marcott et al., 2013); (f) EPICA Dronning Maud Land (EDML)
δ18Oice (EPICA community members, 2006).
In order to perform 17O-excess measurements on water samples at
LSCE, we follow the method described in details in Barkan and Luz (2005). In
short, for each sample, 2 µL of water is injected into a helium flow
purified by passing through a trap immersed in liquid nitrogen. Water vapor
then reacts with CoF3 (producer Sigma-Aldrich) in a nickel tube
heated at 370 ∘C to produce oxygen and hydrofluoric acid, which is
trapped in liquid nitrogen at the outlet of the nickel tube. Oxygen is first
trapped in a molecular sieve tube immersed in liquid nitrogen and then
separated from helium and purified through two cycles of warming
(+30 ∘C) and cooling (-196 ∘C) of the tube with
molecular sieves. The oxygen is finally trapped in a manifold immersed in
liquid helium. The produced oxygen is injected into the mass spectrometer
(MAT 253) and analysed by dual inlet against a reference O2 gas (two runs of 20 measurements).
Every day, at least one home standard is run with the batch of samples to
check the stability of the fluorination line and mass spectrometer. In
addition, a series of three water home standards, whose δ18O
and δ17O values are calibrated on the SMOW–SLAP scale
following Schoenemann et al. (2013), is run at least every month. For this
study, the SMOW–SLAP-calibrated home standards have δ18O
values of -18.64 ‰, -33.56 ‰, and
-54.05 ‰, respectively, hence bracketing the δ18O values of the
measured samples. The comparison of the measured and SMOW–SLAP-calibrated
δ18O and δ17O values then enables the
calibration of
the δ18O and 17O-excess values of the NGRIP samples
following the method described in Schoenemann et al. (2013) and Landais et al.
(2014). The resulting mean uncertainty is 5 ppm (1σ) for the
17O-excess measurements of this study. Note that the use of the MAT
253 mass spectrometer gave more stable results that a Delta V+ instrument
used for previous studies at LSCE (e.g., Landais et al., 2012).
Results: water isotopic records at NGRIP
Our 1518 new measurements of δ18O and d-excess on the NGRIP
ice core cover the time period 14.5 to 60 ka (Fig. 1) and we present
435 duplicate measurements of 17O-excess over the time period
ranging from 9.6 to 20 ka (Fig. 2). δ18O is a qualitative
proxy for local surface temperature. Comparisons between ice core
δ18O data and paleotemperature estimates from borehole
temperature profile inversion and abrupt temperature changes inferred from
isotopic measurements on trapped air showed that the
δ18O–temperature relationship at NGRIP varies from
0.3 ‰ to 0.5 ‰ ∘C-1 during
glacial–interglacial periods (Buizert et al., 2014; Dahl-Jensen et al., 1998;
Kindler et al., 2014).
The second-order parameter d-excess (Dansgaard, 1964) is used in Greenland
ice cores to track past changes in evaporation conditions or shifts in
moisture sources (Johnsen et al., 1989; Masson-Delmotte et al., 2005a).
Evaporation conditions affect the initial vapor d-excess through the impact
of surface humidity and sea surface temperature on kinetic fractionation
(Jouzel et al., 1982). Recent vapor monitoring and modeling studies show
that the d-excess signal of the moisture source can be preserved in polar
vapor and precipitation after transportation towards polar regions (Bonne et
al., 2015; Pfahl and Sodemann, 2014). This signal, however, can be altered
during distillation due to the sensitivity of equilibrium fractionation
coefficients to temperature, leading to alternative definitions using
logarithm formulations for Antarctic ice cores (Uemura et al., 2012; Markle
et al., 2016). Finally, changes in δ18Oseawater also
influence δ18O and d-excess in polar precipitation.
Summarizing, d-excess in Greenland ice cores is a complex tracer:
interpreting its past variations in terms of changes in evaporation
conditions (sea surface temperature or humidity) requires deconvolution of
the effects of glacial–interglacial changes in δ18Oseawater and condensation temperature.
17O-excess provides complementary information to d-excess
(Landais et al., 2008, 2012). At evaporation, d-excess and
17O-excess are both primarily influenced by the balance between
kinetic and equilibrium fractionation, itself driven by relative humidity at
the sea surface. During transport, while d-excess is influenced by
distillation effects during atmospheric cooling, 17O-excess is
largely insensitive to this effect, except at very low temperatures in
Antarctica (Winkler et al., 2012). Conversely, 17O-excess is
affected by recycling or mixing of air masses along the transport path from
low to high latitudes (Risi et al., 2010) and by the range over which
supersaturated conditions occur, itself affected, for instance, by changes in
sea ice extent or temperature along the transport path (Schoenemann et al.,
2014). Because of its logarithmic definition, 17O-excess is not
sensitive to changes in δ18Oseawater given that the
17O-excess of global seawater remains constant with time. As a
consequence, a change in seawater isotopic composition will only be
transmitted to the 17O-excess of the precipitation if the
17O-excess of the evaporated seawater is modified.
As previously reported for the central Greenland GRIP ice core
(Masson-Delmotte et al., 2005b; Jouzel et al., 2005), the NGRIP δ18O and d-excess records exhibit a systematic anticorrelation during
the abrupt Dansgaard–Oeschger (DO) events of the last glacial period and last
deglaciation (Bølling–Allerød and Younger Dryas), with d-excess being
higher during cool Greenland stadials and lower during warm Greenland
interstadials.
The origin of moisture may be different at GRIP and NGRIP. While both sites
are expected to receive most of their moisture from the North Atlantic and
North America (Werner et al., 2001; Landais et al., 2012; Langen and Vinther,
2009) with modulation partly linked to sea ice extent (Rhines et al., 2014),
the northwestern NGRIP site may also receive moisture from the North Pacific
(Langen and Vinther, 2009). Nevertheless, the two sites depict similar
amplitudes of d-excess variations across DO events (Fig. 1). We note that
this contrasts with a slightly lower amplitude (typically by 1 ‰) of
abrupt δ18O changes at NGRIP compared to GRIP.
The fact that d-excess increases (by 3.5±1 ‰) when
δ18O decreases (by 4±1 ‰) during Greenland
stadials relative to interstadials may at least partly reflect the influence
of local temperature changes on d-excess, challenging a simple
interpretation in terms of changes in source conditions. We note one
exception: the Heinrich Stadial 1 cold phase preceding the onset of the
Bølling–Allerød at 14.7 ka. During the time interval 17.5 to 14.7 ka,
the δ18O values measured in the three Greenland ice cores
NGRIP, GRIP, and GISP2 remain almost stable (Fig. 2). Over this period,
δ18O variations are smaller than 1 ‰, i.e., less than
one-fourth of the average amplitude in δ18O changes across DO
events, suggesting no large temperature change in Greenland during this
period. The link between flat δ18O and minimal temperature
variability can be challenged since a mean temperature signal can be masked
by a change in seasonality of moisture source origin on the δ18O record (Boyle et al., 1994; Krinner et al., 1997). However, our
assumption of stable temperature is supported by constant δ15N
of N2 values in the GISP2 and NGRIP ice cores (Buizert et al.,
2014), δ15N of N2 being an alternative
paleothermometry tool in ice cores that is not affected by processes within
the water cycle (Severinghaus and Brook, 1999). In contrast to this almost
stable δ18O signal, d-excess depicts an average
2.2 ‰ increase at 16.1 ka (more than 60 % of the average
amplitude during DO events) with a larger amplitude at GRIP (2.7 ‰ )
than at NGRIP (1.7 ‰) (Fig. 2). In this case, the increase in
d-excess cannot be explained by any Greenland temperature change and
therefore demonstrates a decoupling between cold and stable Greenland
temperatures and changing climatic conditions at lower latitudes during
Heinrich Stadial 1 (see also Supplement).
While the 17O-excess level is similar at the Last Glacial Maximum
(i.e., before 19 ka on Fig. 2) and the early Holocene (40 ppm), it also
shows significant variations during the last deglaciation. Most of these
variations covary with those of δ18O such as the four main
oscillations during the Bølling–Allerød and the onset and end of the
Younger Dryas. They can be interpreted as parallel variations in the
Greenland temperature and lower-latitude climate with a possible contribution
of local temperature through kinetic effects. Again, a major difference
occurs during Heinrich Stadial 1. While the δ18O record is
relatively stable, the 17O-excess profile exhibits a decreasing
trend (strongest between 17.5 and 16.1 ka) before a minimum level is reached
between 16.1 and 14.7 ka. We therefore observe a clear and synchronous signal
in both d-excess and 17O-excess dated around 16.2 ka from
statistical analysis (cf. Sect. 1 in Supplement).
Discussion
The 17O-excess and d-excess transitions at 16.2 ka do not have
any clear counterpart in δ18O (cf. Sect. 2 in Supplement) and no
temperature variation at that time was recorded in the δ15N of
N2 record. We interpret these patterns as illustrating a
reorganization of climatic conditions and/or water cycle at latitudes south
of Greenland. A similar shift in 17O-excess has already been
observed during Heinrich Stadial 4 in the NEEM ice core, while the
δ18O record exhibits a constant low level (Guillevic et al.,
2014). This pattern was also attributed to a change in the water cycle and/or
climate at lower latitudes.
The Greenland water stable isotope records demonstrate a change in the water
cycle and/or climate at lower latitudes at 16.2 ka when Greenland conditions
were relatively stable and cold. This change at low latitudes is confirmed by
the high-resolution atmospheric CH4 concentration record from the
WAIS Divide ice core (Rhodes et al., 2015) (Fig. 2). At 16.2 ka, the
CH4 record indeed exhibits a 30 ppbv peak hypothesized to reflect
more CH4 production in Southern Hemisphere wetlands driven by
wetter conditions due to a southward shift of the tropical rain belts
associated with the ITCZ (Rhodes et al., 2015). The parallel increase in
atmospheric CO2 concentration by 10 ppm in ∼100 years
(Marcott et al., 2013) is understood to result from increased terrestrial
carbon fluxes or enhanced air–sea gas exchange in the Southern Ocean (Bauska
et al., 2014). We also highlight an unusual characteristic of the bipolar
seesaw pattern in Antarctic ice core δ18O records at 16.2 ka.
As observed during all Greenland stadials of the last glacial period,
Antarctic δ18O also increases during Heinrich Stadial 1 (e.g.,
EPICA community members, 2006) through the warming phase of Antarctic
Isotopic Maximum 1. The EPICA Dronning Maud Land (EDML) ice core, drilled in
the Atlantic sector of Antarctica, shows an associated two-step δ18O increase. The first step, marked by a strong increasing trend, is
followed by a change in slope at 16.2 ka. The second step is characterized
by a slower increasing trend from 16.2 to 14.7 ka (EPICA community members,
2006; Stenni et al., 2011) (Fig. 2). The EDML δ18O variations
are expected to be closely connected to changes in AMOC due to the position
of the ice core site on the Atlantic sector of the East Antarctic plateau and
a link between the EDML δ18O record and the low-latitude signal over
Heinrich Stadial 1 has already been suggested by Zhang et al. (2016). For
other Antarctic sites, the change in slope around 16.2 ka is less clear,
probably due to the damping effect of the Southern Ocean or because other
climatic effects linked to atmospheric teleconnections with the tropics
affect the Pacific and Indian sectors of Antarctica (Stenni et al., 2011;
WAIS Divide members, 2013; Buiron et al., 2012). A change in the
teleconnections between West Antarctic climate and tropical regions is also
observed around 16.2 ka (Jones et al., 2018). Summarizing, our synthesis of
ice core records clearly demonstrates a climate shift at 16.2 ka, identified
in proxy records sensitive to shifts in tropical hydrology (CH4),
midlatitude hydrological cycle changes in the Atlantic basin (Greenland
second order isotopic tracers), and in Antarctic climate dynamics in
the Atlantic basin. This suggests some reorganization of water cycle in the
Atlantic region (possibly involving AMOC) related to surface shifts in the
ITCZ at 16.2 ka. This does not appear to affect the high latitudes of the
North Atlantic as Greenland temperatures stay uniformly cold. Uniformly cold
conditions in Greenland are generally observed during Heinrich Stadials of
the last glacial period with temperature and δ18O levels that
are not significantly lower than temperature levels observed during Greenland
stadials (Kindler et al., 2014; Guillevic et al., 2014). Because Greenland is
surrounded by large sea ice during Greenland stadials and Heinrich Stadials
(Hoff et al., 2016), an explanation may be that central Greenland
temperatures are saturated during cold periods so that AMOC modifications
occurring south of the sea ice edge are not significantly influencing
Greenland temperatures. During Heinrich 1, occurring during the last
deglaciation, the situation may be more complicated because of the
CO2 concentration and insolation increases. In this case, the
occurrence of Heinrich Stadial 1 may counteract the increase in Greenland
temperature records induced by CO2 and insolation forcing through
winter cooling driven by AMOC weakening, as suggested by Buizert et
al. (2018).
At low latitudes, an ITCZ shift at 16.2 ka is clearly expressed through a
weak monsoon interval in East Asian speleothem records and through a change in
hydrology in the low-latitude Pacific region, Cariaco Basin, and Brazil
(Partin et al., 2007; Deplazes et al., 2013; Russell et al., 2014; Strikis et
al., 2015). Since we have ruled out a local temperature signal at 16.2 ka in
Greenland, the origin of the Greenland d-excess and 17O-excess
changes around 16.2 ka is also linked to changes in the climate of the
source evaporative regions. When evaporation conditions change, they affect
the proportion of kinetic versus equilibrium fractionation and cause similar
trends in both d-excess and 17O-excess. Both of them indeed
increase when kinetic fractionation is more important, i.e., when relative
humidity decreases or when a change in sea ice modifies the evaporative
conditions (Klein et al., 2015; Kopec et al., 2016). However, d-excess in
the atmospheric vapor is affected by distillation toward higher latitudes
and strongly depends on the source–site temperature gradient, while
17O-excess better preserves the initial fingerprint of the relative
humidity of the evaporative region.
The sequence of phase 1 and phase 2 of Heinrich Stadial 1 identified
in Greenland records and in proxy records of North Atlantic SST, IRD events,
and changes in East Asian hydroclimate. (a) NGRIP (dark blue) and
GRIP (light blue) δ18O records; (b) NGRIP (dark green) and
GRIP (light green) d-excess records; (c) sea surface temperature
(SST) for North Atlantic cores SU 81-18 (Bard et al., 2000) and ODP 161-976
(Martrat et al., 2014); (d) calcite δ18O of Hulu Cave
(China, Zhang et al., 2014); (e)Ca/Sr from site U1308 in
the IRD belt (Hodell et al., 2019) as a signature from strong iceberg
discharges from the Laurentide ice sheet; (f) indications for
Channel River sediment load (blue, sediment load; red, turbidite frequency)
(Toucanne et al., 2010, 2015) as a signature for meltwater input from the European
side. The three red circles indicate plumite layers resulting from outburst
floods on the eastern Canadian margin (Leng et al., 2018), i.e., meltwater
arrival from the North America side in the absence of strong iceberg
discharge. The dashed horizontal line separates the ice core records reported
on the GICC05 timescale from non-ice-core records displayed on their own
timescales.
As a result, the opposing trends observed in d-excess and
17O-excess at 16.2 ka can most probably be explained by an
increase in both the relative humidity and the sea surface temperature of the
evaporative source regions for central and northern Greenland. Despite known
limitations (Winkler et al., 2012; Schoenemann and Steig, 2016), the
classical approach for inferring changes in source relative humidity and
surface temperature from d-excess and 17O-excess in Greenland
(Masson-Delmotte et al., 2005a; Landais et al., 2012) suggests respective
increases of the order of 3 ∘C and 8 % for temperature and
relative humidity of the source evaporative regions, respectively. The larger
d-excess increase at the transition between phase 1 and phase 2 of Heinrich
Stadial 1 observed at GRIP compared to NGRIP is compatible with a larger
proportion of GRIP moisture provided by the midlatitude North Atlantic for
this site. A larger increase in the sea surface temperature of the source of
moisture for GRIP compared to NGRIP would also reduce the source–site
temperature gradient and is fully compatible with the 2 ‰ less
depleted level of δ18O at GRIP, compared to NGRIP, during
phase 2. The increases in both temperature and relative humidity of the
Greenland source regions suggest a more intense evaporative flux from lower
latitudes starting at 16.2 ka. Such features could be explained either by a
local climate signal of evaporative regions or by a southward shift of
evaporative source regions toward warmer and more humid locations. The signal
of source temperature increase is in line with earlier interpretations of
Greenland d-excess changes (Steffensen et al., 2018; Masson-Delmotte et
al., 2005b). The signal of source humidity increase may at least partly
be explained by wetter conditions in the continental North American evaporative
source regions, which are known to partly affect Greenland moisture today in
addition to the main source in the northern Atlantic (Werner et al., 2001; Langen
and Vinther, 2009). This relative humidity signal reconstructed from
Greenland 17O-excess at the transition between phase 1 and phase 2
of Heinrich Stadial 1 coincides with the onset of the “big wet” period in
North American records (Broecker and Putnam, 2012). This transition to a “big
wet” period can be explained by a southward shift of the storm tracks and
polar jet stream over North America during this period (Asmerom et al.,
2010).
We now explore paleoceanographic records to search for a fingerprint of
climate and/or AMOC reorganization at 16.2 ka in the North Atlantic region
and possible implications for our ice core records. Such a comparison of ice
core and marine sediment records appears insightful despite existing
limitations attached to relative chronologies. First, high-resolution proxy
records of surface sea temperature in the East Atlantic, near Europe, depict
a clear warming in the middle of Heinrich Stadial 1 (Bard et al., 2000;
Matrat et al., 2014, Fig. 3). This signal is coherent with our interpretation
of the Greenland d-excess increase at 16.2 ka. In the deep western Atlantic,
no specific feature emerges between phase 1 and phase 2 of Heinrich Stadial 1
from the multi-centennial resolution record of Pa/Th, a proxy for
AMOC strength (McManus et al., 2004).
Heinrich Stadial 1 is associated with at least two major iceberg rafted
debris (IRD) discharges first identified near the Iberian margin (Bard et
al., 2000). They may reflect the impact of changes in ocean conditions
on ice shelf and ice sheet stabilities (Alvarez-Solas et al., 2011).
Alternatively, the iceberg discharges themselves may have affected the AMOC,
which is known to have major impacts on patterns of sea surface temperature,
sea ice, atmospheric circulation, and climate over surrounding continents.
The first IRD phase originated from ice sheet discharges from northern Europe
and Iceland, causing strong reorganizations in deep circulation of the northeast Atlantic
(Stanford et al., 2011; Grousset et al., 2001; Peck et al.,
2006), while the second IRD phase is caused by discharges from the Laurentide
ice sheet. Recent studies (e.g., Hodell et al., 2017; Toucanne et al., 2015)
suggest that all IRD phases occur after 16.2 ka, during Heinrich Stadial 1
phase 2. Before that, Heinrich Stadial phase 1 is associated with a strong
increase in sediment fluxes due to meltwater arrival through terrestrial-terminating ice streams originating from both European and American sides of
the North Atlantic as a response to the beginning of the deglaciation
(Toucanne et al., 2015; Ullman et al., 2015; Leng et al., 2018) (Fig. 3).
During the first slowdown of AMOC during phase 1 of Heinrich Stadial 1, the
associated warming of subsurface water would hence enable the destabilization
of marine ice shelves occurring during phase 2 (Alvarez-Solas et al., 2011;
Marcott et al., 2011). This second phase of Heinrich Stadial 1 is also
associated with extensive sea ice production, south of Greenland
(Hillaire-Marcel and de Vernal, 2008). The increase in North Atlantic sea ice
extent and major iceberg discharges during the second phase of Heinrich
Stadial 1 are coherent with a southward shift of the evaporative region,
providing moisture to Greenland supported by d-excess data, and a southward
shift of tropical rain belts (Chiang and Bitz, 2005), affecting Southern
Hemisphere CH4 sources (Rhodes et al., 2015).
Conclusions
Combined measurements of d-excess and 17O-excess along the NGRIP
ice core demonstrate a decoupling between a cold and stable Greenland climate
and changes in hydroclimate at lower latitudes during the Heinrich Stadial 1,
also referred to as the “mystery interval” (Denton et al., 2006). These new
measurements hence confirm the previous studies of Zhang et al. (2014, 2016)
and Rhodes et al. (2015). While Greenland temperature remains mostly stable
from 20 to 14.7 ka, a large-scale climatic reorganization takes place at
16.2 ka, associated with warmer and wetter conditions at the location of
Greenland moisture sources. Based on a coherent temporal framework linking
the different ice core records, we show that this event coincides with
changes in the characteristics of the bipolar seesaw pattern, as observed in
the Atlantic sector of Antarctica, and has a fingerprint in global
atmospheric composition through sharp changes in atmospheric CO2
and CH4 concentrations.
Based on these new ice core records and the comparison with marine and
terrestrial records, we propose the following sequence of events during the
last deglaciation. First, the initiation of Heinrich Stadial 1 occurs at
17.5 ka or earlier, with meltwater arrival from the terrestrial-terminating
ice streams synchronous with a decrease in the North Atlantic sea surface
temperature offshore of Europe, a first AMOC slowdown, drier conditions in
North America, and an increase in Antarctic temperature as well as in
atmospheric CO2 and CH4 concentrations. No fingerprint of
this first phase of Heinrich Stadial 1 is identified in Greenland water
stable isotope records; δ18O (and thus local temperature),
17O-excess, and d-excess remain stable. A possible explanation for
such stability is that the high-latitude warming induced by the increase in
the summer insolation at high latitudes over the beginning of the deglaciation
is counterbalanced in Greenland by regional changes in, e.g., increased albedo
due to sea ice extent or reduced transport of heat by the atmospheric
circulation towards central Greenland, which both can result from a reduced
AMOC strength. The global event occurring at 16.2 ka marks the onset of the
second phase of Heinrich Stadial 1. It is associated with (i) strong iceberg
discharges due to dynamical instability of the Laurentide ice sheet, probably
induced by the accumulation of subsurface ocean heat due to a slowdown of
AMOC during phase 1, (ii) a widespread reorganization of the atmospheric
water cycle in the Atlantic region, with significant changes in d-excess
and 17O-excess in Greenland, (iii) the initiation of
a weak monsoon interval in East Asia, and (iv) the transition from a “big dry”
episode to a “big wet” episode in North America. We note that this
sequence of events within Heinrich Stadial 1 is invisible in all available
Greenland temperature proxy records, which only display an abrupt warming at
the onset of the Bølling–Allerød (14.7 ka).
Attached to a bipolar synchronized chronological framework, our new ice core
data provide a unique benchmark to test the ability of Earth system models
(ESMs) to correctly resolve the mechanisms occurring at sub-millennial scale
during the last deglaciation. In particular, our proposed sequence of events
is useful to evaluate ESM capabilities in reproducing the relationships
between meltwater fluxes, the state of the North Atlantic ocean circulation,
the Laurentide ice sheet instability, changes at the moisture sources of
Greenland ice cores, the response of hydroclimate at low and high latitudes,
and the net quantitative effects on global methane and carbon
budgets.
The new datasets are given in the Supplement below.
The supplement related to this article is available online at: https://doi.org/10.5194/cp-14-1405-2018-supplement.
AL, EC and VMD designed the study. BM, FP and AL performed the measurements.
AL, EC, ST, RR and TP worked on interpretation of the results. AL wrote the first draft and all authors contributed to the writing of the final manuscript.
The authors declare that they have no conflict of
interest.
Acknowledgements
The research leading to these results has received funding from the European
Research Council under the European Union's Seventh Framework Programme
(FP7, 2007–2013) RC agreement number 306045. Emilie Capron is funded by the
European Union's Seventh Framework Programme for research and innovation
under Marie Skłodowska-Curie grant agreement no. 600207.
Rachael H. Rhodes received funding from a
European Commission Horizon 2020 Marie Skłodowska-Curie Individual Fellowship
(no. 658120, SEADOG). We thank Myriam Guillevic and Jean Jouzel for very
useful discussions as well as Christo Buizert and an anonymous reviewer for
very useful comments. Edited by: Elizabeth
Thomas Reviewed by: Christo Buizert and one anonymous referee
ReferencesÁlvarez-Solas, J., Montoya, M., Ritz, C., Ramstein, G., Charbit, S.,
Dumas, C., Nisancioglu, K., Dokken, T., and Ganopolski, A.: Heinrich event 1:
an example of dynamical ice-sheet reaction to oceanic changes, Clim. Past, 7,
1297–1306, 10.5194/cp-7-1297-2011, 2011.
Andersen, K. K., Svensson, A., Johnsen, S. J., Rasmussen, S. O., Bigler, M.,
Roethlisberger, R., Ruth, U., Siggaard-Andersen, M. L., Steffensen, J. P.,
Dahl-Jensen, D., Vinther, B. M., and Clausen, H. B., The Greenland Ice Core
Chronology 2005, 15-42 ka. Part 1: constructing the time scale, Quaternary
Sci. Rev., 25, 3246–3257, 2006
Asmerom, Y., Polyak, V. J., and Burns, S. J.: Variable winter moisture in the
southwestern United States linked to rapid glacial climate shifts, Nat.
Geosci., 3, 114–117, 2010.
Bard, E., Rostek, F., Turon, J. L., and Gendreau, S.: Hydrological Impact of
Heinrich Events in the Subtropical Northeast Atlantic, Science, 289,
1321–1324, 2000Barkan, E. and Luz, B.: High precision measurements of 17O/16O
and 18O/16O ratios in H2O, Rapid Commun., 19,
3737–3742, 2005.
Bauska, T. K., Baggenstos, D., Brook, E. J., Mix, A. C., Marcott, S. A., and
Petrenko, V.: Carbon isotopes characterize rapid changes in atmospheric
carbon dioxide during the last deglaciation, P. Natl. Acad. Sci. USA, 113,
3465–3470, 2016.Bazin, L., Landais, A., Lemieux-Dudon, B., Toyé Mahamadou Kele, H.,
Veres, D., Parrenin, F., Martinerie, P., Ritz, C., Capron, E., Lipenkov, V.,
Loutre, M.-F., Raynaud, D., Vinther, B., Svensson, A., Rasmussen, S. O.,
Severi, M., Blunier, T., Leuenberger, M., Fischer, H., Masson-Delmotte, V.,
Chappellaz, J., and Wolff, E.: An optimized multi-proxy, multi-site Antarctic
ice and gas orbital chronology (AICC2012): 120–800 ka, Clim. Past, 9,
1715–1731, 10.5194/cp-9-1715-2013, 2013.
Bonne, J.-L., Steen-Larsen, H. C., Risi, C., Werner, M., Sodemann, H.,
Lacour, J.-L., Fettweis, X., Cesana, G., Delmotte, M., Cattani, O.,
Vallelonga, P., Kjaer, H. A., Clerbaux, C., Sveinbjornsdottir, A. E., and
Masson-Delmotte, V.: The summer 2012 Greenland heat wave: In situ and remote
sensing observations of water vapor isotopic composition during an
atmospheric river event, J. Geophys. Res.-Atmos., 120, 2970–2989, 2015.Boyle, E. A.: Cool tropical temperatures shift the global δ18O-T relationship: An explanation for the ice core δ18O-borehole thermometry conflict?, Geophys. Res. Lett., 24, 273–276,
1994.
Broecker, W. and Putnam, A. E.: How did the hydrologic cycle respond to the
two-phase mystery interval?, Quaternary Sci. Rev., 57, 17–25, 2012.
Buiron, D., Stenni, B., Chappellaz, J., Landais, A., Baumgartner, M.,
Bonazza, M., Capron, E., Frezzotti, M., Kageyama, M., Lemieux-Dudon, B.,
Masson-Delmotte, V., Parrenin, F., Schilt, A., Selmo, E., Severi, M.,
Swingedouw, D., and Udisti, R.: Regional imprints of millennial variability
during the MIS 3 period around Antarctica, Quaternary Sci. Rev., 48, 99–112,
2012.
Buizert, C., Gkinis, V., Severinghaus, J. P., He, F., Lecavalier, B. S.,
Kindler, P., Leuenberger, M., Carlson, A. E., Vinther, B., Masson-Delmotte,
V., White, J. W. C., Liu, Z., Otto-Bliesner, B., and Brook, E. J.: Greenland
temperature response to climate forcing during the last deglaciation,
Science, 345, 1177–1180, 2014.
Buizert, C., Keisling, B. A., Box, J. E., He, F., Carlson, A. E., Sinclair,
G., and DeConto, R. M.: Greenland-Wide Seasonal Temperatures During the Last
Deglaciation, Geophys. Res. Lett., 45, 1905–1914, 2018.
Chiang, J. C. H. and Bitz, C. M.: Influence of high latitude ice cover on the
marine Intertropical Convergence Zone, Clim. Dynam., 25, 477–496, 2005.
Clark, P. U., Shakun, J. D., Baker, P. A., Bartlein, P. J., Brewer, S.,
Brook, E., Carlson, A. E., Cheng, H., Kaufman, D. S., Liu, Z., Marchitto, T.
M., Mix, A. C., Morrill, C., Otto-bliesner, B. L., Pahnke, K., Russell, J.
M., Whitlock, C., Adkins, J. F., Blois, J. L., Clark, J., Colman, S. M.,
Curry, W. B., Flower, B. P., He, F., Johnson, T. C., Lynch-Stieglitz, J.,
Markgraf, V., Mcmanus, J., Mitrovica, J. X., Moreno, P. I., and Williams, J.
W.: Global climate evolution during the last deglaciation, P. Natl. Acad.
Sci. USA, 109, E1134–E1142, 2012.
Dahl-Jensen, D., Mosegaard, K., Gundestrup, N., Clow, G. D., Johnsen, S. J.,
Hansen, A. W., and Balling, N.: Past Temperatures Directly from the Greenland
Ice Sheet, Science, 282, 268–271, 1998.
Dansgaard, W.: Stable isotopes in precipitation, Tellus, 16, 436–468, 1964.
Denton, G. H., Broecker, W. S., and Alley, R. B.: The mystery interval 17.5
to 14.5 kyrs ago, PAGES news, 14,
14–16, 2006.
Denton, G. H., Anderson, R. F., Toggweiler, J. R., Edwards, R. L.,Schaefer,
J. M., and Putnam, A. E.: The Last Glacial Termination, Science, 328,
1652–1656, 2010.
Deplazes, G., Luckge, A., Peterson, L. C., Timmermann, A., Hamann, Y.,
Hughen, K. A., Rohl, U., Laj, C., Cane, M. A., Sigman, D. M., and Haug, G.
H.: Links between tropical rainfall and North Atlantic climate during the
last glacial period, Nat. Geosci., 6, 213–217, 2013.
Epica community members: One-to-one coupling of glacial climate variability
in Greenland and Antarctica, Nature, 444, 195–198, 2006.
Epstein, S., Buchsbaum, R., Lowenstam, H. A., and Urey, H.: Revised
carbonate-water isotopic temperature scale, GSA bulletin, 64, 1315–1325, 1953.
Grootes, P. M., Stuiver, M., White, J. W. C., Johnsen, S., and Jouzel, J.:
Comparison of oxygen isotope records from the GISP2 and GRIP Greenland ice
cores, Nature, 366, 552–554, 1993.
Grousset, F.: Les changements abrupts du climat depuis 60 000 ans/Abrupt
climatic changes over the last 60,000 years, Quaternaire, 12, 203–211, 2001.Guillevic, M., Bazin, L., Landais, A., Stowasser, C., Masson-Delmotte, V.,
Blunier, T., Eynaud, F., Falourd, S., Michel, E., Minster, B., Popp, T.,
Prié, F., and Vinther, B. M.: Evidence for a three-phase sequence during
Heinrich Stadial 4 using a multiproxy approach based on Greenland ice core
records, Clim. Past, 10, 2115–2133, 10.5194/cp-10-2115-2014,
2014.
Hillaire-Marcel, C. and De Vernal, A.: Stable isotope clue to episodic sea
ice formation in the glacial North Atlantic, Earth Plan. Sc. Let., 268,
143–150, 2008.Hoff, U., Rasmussen, T. L., Stein, R., Ezat, M. M., and Fahl, K.: Sea ice and
millennial-scale climate variability in the Nordic seas 90-kyr ago to
present, Nat. Commun., 7, 12247, 10.1038/ncomms12247, 2016.Hodell, D. A., Nicholl, J. A., Bontognali, T. R. R., Danino, S., Dorador, J.,
Dowdeswell, J. A., Einsle, J., Kuhlmann, H., Martrat, B., Mleneck-vautravers,
M. J., Rodríguez-tovar, F. J., and Röhl, U.: Anatomy of Heinrich
Layer 1 and its role in the last deglaciation, Paleoceanogr. Paleocl., 32,
10.1002/2016PA003028, 2017.
Johnsen, S. J., Dansgaard, W., and White, J. W. C.: The origin of Arctic
precipitation under present and glacial conditions, Tellus B, 41, 452–468,
1989.
Jones, T. R., Roberts, W. H. G., Steig, E. J., Cuffey, K. M., Markle, B. R.,
and White, J. W. C.: Southern Hemisphere climate variability forced by
Northern Hemisphere ice-sheet topography, Nature, 554, 351–355, 2018.
Jouzel, J., Merlivat, L., and Lorius, C.: Deuterium excess in an East
Antarctic ice core suggests higher relative humidity at the oceanic surface
during the last glacial maximum, Nature, 299, 688–691, 1982.
Jouzel, J., Masson-Delmotte, V., Stiévenard, M., and Landais, A.: Rapid
deuterium-excess changes in Greenland ice cores?: a link between the ocean
and the atmosphere, C. R. Geosci., 337, 957–969, 2005.Kindler, P., Guillevic, M., Baumgartner, M., Schwander, J., Landais, A., and
Leuenberger, M.: Temperature reconstruction from 10 to 120 kyr b2k from the
NGRIP ice core, Clim. Past, 10, 887–902,
10.5194/cp-10-887-2014, 2014.
Klein, E. S., Cherry, J. E., Young, J., Noone, D., Leffler, A. J., and
Welker, J. M.: Arctic cyclone water vapor isotopes support past sea ice
retreat recorded in Greenland ice, Sci. Rep., 5, 10295–10300, 2015.
Kopec, B. G., Feng, X., Michel, F. A., and Posmentier, E. S.: Influence of
sea ice on Arctic precipitation, P. Natl. Acad. Sci. USA, 113, 46–51, 2016.Krinner, G., Genthon, C., and Jouzel, J.: GCM analysis of local influences on
ice core δ signals, Geophys. Res. Lett., 24, 2825–2828, 1997.Landais, A., Barkan, E., and Luz, B.: Record of δ18O and
17O-excess in ice from Vostok Antarctica during the last 150,000
years, Geophys. Res. Lett., 35, L02709, 10.1029/2007GL032096, 2008.
Landais, A., Steen-Larsen, H. C., Guillevic, M., Masson-Delmotte, V., Vinther
and B., and Winkler, R.: Triple isotopic composition of oxygen in surface
snow and water vapor at NEEM (Greenland), Geochim. Cosmochim. Ac., 77,
304–316, 2012.
Landais, A., Winkler, R., and Prié, F.: Triple Isotopic Composition of
Oxygen in Water from Ice Cores, application note 30287, Thermo, 2014.
Langen, P. L. and Vinther, B. M.: Response in atmospheric circulation and
sources of Greenland precipitation to glacial boundary conditions, Clim.
Dynam., 32, 1035–1054, 2009.Leng, W., von Dobeneck, T., Bergmann, F., Just, J., Mulitza, S., Chiessi, C.,
St-Onge, G., and Piper, D.: Sedimentary and rock magnetic signatures and
event scenarios of deglacial outburst floods from the Laurentian Channel Ice
Stream, Quaternary Sci. Rev., 186, 27–46,
10.1016/j.quascirev.2018.01.016, 2018.Marcott, S. A., Clark, P. U., Padman, L., Klinkhammer, G. P., Springer, S.
R., Liu, Z., Otto-Bliesner, B. L., Carlson, A. E., Ungerer, A., Padman, J.,
He, F., Cheng, J., and Schmittner, A.: Ice-shelf collapse from subsurface
warming as a trigger for Heinrich events, P. Natl. Acad. Sci.
USA, 108, 13415–13419, 10.1073/pnas.1104772108, 2011.
Marcott, S. A, Bauska, T. K., Buizert, C., Steig, E. J., Rosen, J. L.,
Cuffey, K. M., Fudge, T. J., Severinghaus, J. P., Ahn, J., Kalk, M. L.,
McConnell, J. R., Sowers, T., Taylor, K. C., White, J. W. C., and Brook, E.
J.: Centennial-scale changes in the global carbon cycle during the last
deglaciation, Nature, 514, 616–619, 2013.
Markle, B. R., Steig, E. J., Buizert, C., Schoenemann, S. W., Bitz, C. M.,
Fudge, T. J., Pedro, J. B., Ding, Q., Jones, T. R., White, J. W. C., and
Sowers, T.: Global atmospheric teleconnections during Dansgaard–Oeschger
events, Nat. Geosci., 10, 36–40, 2016.
Martrat, B., Jimenez-Amat, P., Zahn, R., and Grimalt, J. O.: Similarities and
dissimilarities between the last two deglaciations and interglaciations in
the North Atlantic region, Quateranry Sci. Rev., 99, 122–134, 2014.
Masson-Delmotte, V., Landais, A., Stievenard, M., Cattani, O., Falourd, S.,
Jouzel, J., Johnsen, S. J. Dahl-Jensen, D., Sveinsbjornsdottir, A., White, J.
W. C., Popp, T., and Fischer, H.: Holocene climatic changes in Greenland:
Different deuterium excess signals at Greenland Ice Core Project (GRIP) and
NorthGRIP, J. Geophys. Res.-Atmos., 110, 1–13, 2005a.
Masson-Delmotte, V., Jouzel, J., Landais, A., Stievenard, M., Johnsen, S.J.,
White, J. W. C., Werner, M., Sveinbjornsdottir, A., and Fuhrer, K.: GRIP
deuterium excess reveals rapid and orbital-scale changes in Greenland
moisture origin, Science, 309, 118–121, 2005b.
McManus, J. F., Francois, R., Gherardi, J.-M., Keigwin, L. D., and
Brown-Leger, S.: Collapse and rapid resumption of Atlantic meridional
circulation linked to deglacial climate changes, Nature, 428, 834–837, 2004.
NGRIP community members: High-resolution climate record of Northern
Hemisphere climate extending into the last interglacial period, Nature, 431,
147–151, 2004.
Partin, J. W., Cobb, K. M., Adkins, J. F., Clark, B., and Fernandez, D. P.:
Millennial scale trends in west Pacific warm pool hydrology since the Last
Glacial Maximum, Nature, 449, 452–455, 2007.
Peck, V. L., Hall, I. R., Zahn, R., Elderfield, H., and Grousset, F.: High
resolution evidence for linkages between NW European ice sheet instability
and Atlantic Meridional Overturning Circulation, Earth Plan. Sc. Let., 243,
476–488, 2006.Pfahl, S. and Sodemann, H.: What controls deuterium excess in global
precipitation?, Clim. Past, 10, 771–781,
10.5194/cp-10-771-2014, 2014.
Rasmussen, S. O., Andersen, K. K., Svensson, A. M., Steffensen, J. P.,
Vinther, B. M., Clausen, H. B., Siggaard-Andersen, M. L., Johnsen, S. J.,
Larsen, L. B., Dahl-Jensen, D., Bigler, M., Roethlisberger, R., Fischer, H.,
Goto-Azuma, K., Hansson, M. E., and Ruth, U.: A new Greenland ice core
chronology for the last glacial termination, J. Geophys. Res.-Atmos., 111,
1–16, 2006.
Rasmussen, S. O., Seierstad, I. K., Andersen, K. K., Bigler, M., and Johnsen,
S. J.: Synchronization of the NGRIP , GRIP , and GISP2 ice cores across MIS 2
and palaeoclimatic implications, Quaternary Sci. Res., 27, 18–28, 2008.
Rhines, A. and Huybers, P. J.: Sea ice and dynamical controls on
preindustrial and last glacial maximum accumulation in central Greenland, J.
Climate, 27, 8902–8917, 2014.
Rhodes, R. H., Brook, E. J., Chiang, J. C. H., Blunier, T., Maselli, O. J.,
McConnell, J., Romanini, D., and Severinghaus, J. P.: Enhanced tropical
methane production in response to iceberg discharge in the North Atlantic,
Science, 348, 1016–1019, 2015.Risi, C., Landais, A., Bony, S., Jouzel, J., Masson-Delmotte, V., and Vimeux,
F.: Understanding the 17O excess glacial-interglacial variations in
Vostok precipitation, J. Geophys. Res.-Atmos., 115, D10112, 10.1029/2008JD011535, 2010.
Russell, J., Vogel, H., Konecky, B.L., Bijaksana, S., Huang, Y., Melles, M.,
Wattrus, N., Costa, K., and King, J.W.: Glacial forcing of central Indonesian
hydroclimate since 60,000 y BP, P. Natl Acad. Sci. USA, 111, 5100–5105,
2014.Schoenemann, S. W., Schauer, S., and Steig, E. J.: Measurement of SLAP2 and
GISP-δ17O and proposed VSMOW-SLAP normalization for
δ17O and 17Oexcess, Rapid Commun. Mass Sp., 27, 582–590,
2013.Schoenemann, S. W., Steig, E. J., Ding, Q., Markle, B. R., and Schauer, A.
J.: Triple water-isotopologue record from WAIS Divide, Antarctica: Controls
on glacial-interglacial changes in 17O-excess of precipitation, J.
Geophys. Res.-Atmos., 119, 8741–8763, 2014.Schoenemann, S. W. and Steig, E. J.: Seasonal and spatial variations of
17O-excess and d-excess in Antarctic precipitation: Insights from
an intermediate complexity isotope model, J. Geophys. Res.-Atmos., 121,
11211–215247, 2016.
Severinghaus, J. P. and Brook, E.: Abrupt climate change at the end of the
last glacial period inferred from trapped air in polar ice, Science, 286,
930–934, 1999.
Stanford, D., Rohling, E. J., Bacon, S., Roberts, A. P., Grousset, F. E., and
Bolshaw, M.: A new concept for the paleoceanographic evolution of Heinrich
event 1 in the North Atlantic, Quaternary Sci. Rev., 30, 1047–1066, 2011.
Steffensen, J. P., Andersen, K. K., Bigler, M., Clausen, H. B., Dahl-Jensen,
D., Fischer, H., Goto-Azuma, K., Hansson, M., Johnsen, S. J., Jouzel, J.,
Masson-Delmotte, V., Popp, T., Rasmussen, S. O., Rothlisberger, R., Ruth, U.,
Stauffer, B., Siggaard-Andersen, M.-L., Sveinbjörnsdóttir, A. E.,
Svensson, A., and White, J. W. C.: High-Resolution Greenland Ice Core Data
Show Abrupt Climate Change Happens in Few Years, Science, 321, 680–684, 2008.
Stenni, B., Buiron, D., Frezzotti, M., Albani, S., Barbante, C., Bard, E.,
Barnola, J. M., Baroni, M., Baumgartner, M., Bonazza, M., Capron, E.,
Castellano, E., Chappellaz, J., Delmonte, B., Falourd, S., Genoni, L.,
Iacumin, P., Jouzel, J., Kipfstuhl, S., Landais, A., Lemieux-Dudon, B.,
Maggi, V., Masson-Delmotte, V., Mazzola, C., Minster, B., Montagnat, M.,
Mulvaney, R., Narcisi, B., Oerter, H., Parrenin, F., Petit, J. R., Ritz, C.,
Scarchilli, C., Schilt, A., Schüpbach, S., Schwander, J., Selmo, E.,
Severi, M., Stocker, T. F., and Udisti, R.: Expression of the bipolar see-saw
in Antarctic climate records during the last deglaciation, Nat. Geosci., 4,
46–49, 2011.
Stríkis, N. M., Chiessi, C. M., Cruz, F. W., Vuille, M., Cheng, H., De
Souza Barreto, E. A., Mollenhauer, G., Kasten, S., Karmann, I., Edwards, R.
L., Bernal, J. P., and Sales, H. D. R.: Timing and structure of Mega-SACZ
events during Heinrich Stadial 1, Geophys. Res. Lett., 42, 5477–5484, 2015.Svensson, A., Andersen, K. K., Bigler, M., Clausen, H. B., Dahl-Jensen, D.,
Davies, S. M., Johnsen, S. J., Muscheler, R., Parrenin, F., Rasmussen, S. O.,
Röthlisberger, R., Seierstad, I., Steffensen, J. P., and Vinther, B. M.:
A 60 000 year Greenland stratigraphic ice core chronology, Clim. Past, 4,
47–57, 10.5194/cp-4-47-2008, 2008.
Toucanne, S., Zaragosi, S., Bourillet, J.-F., Marieu, V., Cremer, M.,
Kageyama, M., Van Vliet-Lanoë, B., Eynaud, F., Turon, J.-L., and Gibbard,
P.-L.: The first estimation of Fleuve Manche palaeoriver discharge during the
last deglaciation: Evidence for Fennoscandian ice sheet meltwater flow in the
English Channel ca 20–18 ka ago, Earth Planet. Sc. Lett., 290, 459–473,
2010.
Toucanne, S., Soulet, G., Freslon, N., Silva Jacinto, R., Dennielou, B.,
Zaragosi, S., Eynaud, F., Bourillet, J.-F., and Bayon, G.: Millennial-scale
fluctuations of the European Ice Sheet at the end of the last glacial, and
their potential impact on global climate, Quaternary Sci. Rev., 123,
113–133, 2015.Uemura, R., Masson-Delmotte, V., Jouzel, J., Landais, A., Motoyama, H., and
Stenni, B.: Ranges of moisture-source temperature estimated from Antarctic
ice cores stable isotope records over glacial-interglacial cycles, Clim.
Past, 8, 1109–1125, 10.5194/cp-8-1109-2012, 2012.
Ullman, D. J., Carlson, A. E., Anslow, F. S., LeGrande, A. N., and Licciardi,
J. M.: Laurentide ice-sheet instability during the last deglaciation, Nat.
Geosci., 8, 534–537, 2015.
Vaughn, B., H., White, J. W. C., Delmotte, M., Trolier, M., Cattani, O., and
Stievenard, M.: An automated system for hydrogen isotope analysis of water,
Chemical Geology Including Isotope Geoscience, 152, 309–319, 1998.Veres, D., Bazin, L., Landais, A., Toyé Mahamadou Kele, H.,
Lemieux-Dudon, B., Parrenin, F., Martinerie, P., Blayo, E., Blunier, T.,
Capron, E., Chappellaz, J., Rasmussen, S. O., Severi, M., Svensson, A.,
Vinther, B., and Wolff, E. W.: The Antarctic ice core chronology (AICC2012):
an optimized multi-parameter and multi-site dating approach for the last 120
thousand years, Clim. Past, 9, 1733–1748,
10.5194/cp-9-1733-2013, 2013.Wais Divide Members: Onset of deglacial warming in West Antarctica driven by
local orbital forcing Nature, 500, 440–444, 2013.
Werner, M., Heimann, M., and Hoffmann, G.: Isotopic composition and origin of
polar precipitation in present and glacial climate simulations, Tellus B,
53, 53–71, 2001.Winkler, R., Landais, A., Sodemann, H., Dümbgen, L., Prié, F.,
Masson-Delmotte, V., Stenni, B., and Jouzel, J.: Deglaciation records of
17O-excess in East Antarctica: reliable reconstruction of oceanic normalized
relative humidity from coastal sites, Clim. Past, 8, 1–16,
10.5194/cp-8-1-2012, 2012.
Zhang, W., Wu, J., Wang, Y., Wang, Y., Cheng, H., Kong, X., and Duan, F. A:
detailed East Asian monsoon history surrounding the “Mystery Interval”
derived from three Chinese speleothem records, Quateranry Res., 82, 154–163,
2014.
Zhang, H., Griffiths, M. L., Huang, J., Cai, Y., Wang, C., Zhang, F., Cheng,
H., Ning, Y., Hu, C., and Xie, S.: Antarctic link with East Asian summer
monsoon variability during the Heinrich Stadial–Bølling interstadial
transition, Earth Planet. Sc. Lett., 453, 243–251, 2016.