Sometime during the middle to late Holocene (8.2 ka to ∼ 1850–1900 CE), the Greenland Ice Sheet (GrIS) was smaller than its current
configuration. Determining the exact dimensions of the Holocene ice-sheet
minimum and the duration that the ice margin rested inboard of its current
position remains challenging. Contemporary retreat of the GrIS from its
historical maximum extent in southwestern Greenland is exposing a landscape
that holds clues regarding the configuration and timing of past ice-sheet
minima. To quantify the duration of the time the GrIS margin was near its
modern extent we develop a new technique for Greenland that utilizes in situ
cosmogenic 10Be–14C–26Al in bedrock samples that have become
ice-free only in the last few decades due to the retreating ice-sheet margin at
Kangiata Nunaata Sermia (n=12 sites, 36 measurements; KNS), southwest Greenland. To maximize
the utility of this approach, we refine the deglaciation history of the region
with stand-alone 10Be measurements (n=49) and traditional 14C ages
from sedimentary deposits contained in proglacial–threshold lakes. We combine
our reconstructed ice-margin history in the KNS region with additional
geologic records from southwestern Greenland and recent model simulations of
GrIS change to constrain the timing of the GrIS minimum in southwest
Greenland and the magnitude of Holocene inland GrIS retreat, as well as to explore the
regional climate history influencing Holocene ice-sheet behavior. Our
10Be–14C–26Al measurements reveal that (1) KNS retreated behind
its modern margin just before 10 ka, but it likely stabilized near the
present GrIS margin for several thousand years before retreating farther
inland, and (2) pre-Holocene 10Be detected in several of our sample sites
is most easily explained by several thousand years of surface exposure during
the last interglaciation. Moreover, our new results indicate that the minimum
extent of the GrIS likely occurred after ∼5ka, and the GrIS
margin may have approached its eventual historical maximum extent as early as
∼2ka. Recent simulations of GrIS change are able to match the
geologic record of ice-sheet change in regions dominated by surface mass
balance, but they produce a poorer model–data fit in areas influenced by oceanic
and dynamic processes. Simulations that achieve the best model–data fit
suggest that inland retreat of the ice margin driven by early to middle
Holocene warmth may have been mitigated by increased precipitation. Triple
10Be–14C–26Al measurements in recently deglaciated bedrock
provide a new tool to help decipher the duration of smaller-than-present ice
over multiple timescales. Modern retreat of the GrIS margin in southwest
Greenland is revealing a bedrock landscape that was also exposed during the
migration of the GrIS margin towards its Holocene minimum extent, but it has yet
to tap into a landscape that remained ice-covered throughout the entire
Holocene.
Introduction
The Greenland Ice Sheet (GrIS) has expanded and contracted repeatedly
throughout the Quaternary. During glaciations the GrIS margin extends onto the
continental shelf, whereas during interglaciations, the dimensions of the GrIS
are often similar to or smaller than today (de Vernal and Hillaire-Marcel,
2008; Hatfield et al., 2016; Knutz et al., 2019). Direct evidence of former
GrIS maxima is found in offshore sedimentary deposits (e.g., Ó Cofaigh
et al., 2013; Knutz et al., 2019), and the pattern of retreat from the most
recent ice-sheet maximum can be reconstructed in detail through a combination
of well-dated marine and terrestrial sedimentary archives (Bennike and Bjork,
2002; Funder et al., 2011; Kelley et al., 2013; Hogan et al., 2016; Jennings
et a., 2017; Young et al., 2020a). Reconstructing the size and timing of
ice-sheet minima, however, is extremely challenging because terrestrial
evidence relating to ice-sheet minima has been overrun and destroyed by
subsequent glacier re-expansion or resides in a largely inaccessible
environment beneath modern glacier footprints. In place of direct terrestrial
evidence, sediment-based proxy records contained in offshore depocenters are used to infer the dimensions and timing of paleo-GrIS minima (Colville
et al., 2011; Reyes et al., 2014; Bierman et al., 2016; Hatfield et al.,
2016). These sediment-based approaches are not able to provide direct
constraints on the magnitude or timing of GrIS minima, but they have the
advantage of generally providing continuous records of inferred ice-sheet
change.
Cosmogenic isotope measurements from recently deglaciated bedrock surfaces or
those still residing under ice provide key constraints on the timing and
magnitude of glacier and ice-sheet minima (e.g., Goehring et al., 2011;
Schaefer et al., 2016; Pendleton et al., 2019). These bedrock surfaces serve
as fixed benchmark locations, and nuclide accumulation can only occur under
extremely thin ice (e.g., in situ 14C) or, more commonly, in the
absence of ice cover when surfaces are exposed to the atmosphere (i.e., a
direct ice-margin constraint). The primary caveat of this method, however, is
that measured nuclide inventories have non-unique solutions and only provide a
measure of integrated surface exposure and burial. Moreover, drilling through
extant glaciers and ice sheets to bedrock is logistically challenging,
expensive, and can only be done after lengthy site consideration (e.g.,
Spector et al., 2018). Nonetheless, groundbreaking measurements of cosmogenic
in situ 10Be and 26Al in bedrock beneath the GISP2 borehole
revealed that the GrIS likely disappeared on several occasions during the
Pleistocene (Schaefer et al., 2016).
Contemporary retreat of the GrIS margin from its historical maximum extent is
exposing a fresh bedrock landscape, and inventories of cosmogenic nuclides in
this newly exposed bedrock can provide clues to past ice-sheet minima without
having to drill through ice. Abundant geological evidence reveals that
sometime during the middle Holocene, the GrIS was slightly smaller than today
(e.g., Weidick et al., 1990; Long et al., 2011; Lecavalier et al., 2014;
Larsen et al., 2015; Young and Briner, 2015; Lesnek et al., 2020). The
mid-Holocene minimum was forced by regional temperatures that were likely as
warm or warmer than today, and elucidating the behavior of the GrIS during this
interval can provide key insights into GrIS behavior in a warming
world. Bedrock emerging today from beneath the GrIS margin was potentially ice-free during the middle Holocene, and cosmogenic nuclides in these surfaces can
constrain the magnitude and duration of inland GrIS retreat.
Here, we present in situ cosmogenic
10Be–14C–26Al measurements from recently exposed
bedrock surfaces (n=12 sites) in the Kangiata Nunaata Sermia (KNS)
forefield, southwestern Greenland (Figs. 1 and 2). Triple
10Be–14C–26Al measurements have, to the best of
our knowledge, rarely been made (e.g., Miller et al., 2006; Briner et al.,
2014) and have not been utilized in any systematic fashion in recently
deglaciated environments. To aid interpretation of our
10Be–14C–26Al measurements, we refine the early
Holocene deglaciation history of the landscape immediately outboard of the
historical GrIS maximum extent and constrain when the GrIS retreated inboard
of its present position through a combination of stand-alone 10Be
measurements (n=49) and traditional 14C-dated sediment sequences
from proglacial–threshold lakes. We combine our new results with previously
published records of deglaciation in southwestern Greenland to estimate when
the GrIS was behind its present position and reached its minimum extent. We
compare geologic records of ice-sheet change to recent model simulations of
Holocene GrIS change to further assess the timing and magnitude of
mid-Holocene GrIS retreat.
Settings and methodsOverview
The study region is characterized by mountainous terrain dissected by a dense
fjord network in which KNS resides (Fig. 1). Bedrock in the region consists
primarily of Archean gneiss (Henriksen et al., 2000). Decades of research have
resulted in a robust record of regional deglaciation. Minimum-limiting
radiocarbon ages and 10Be ages reveal that initial coastal
deglaciation occurred at ∼11.2–10.7 ka and the inner fjord
region was ice-free by ∼10.5–10.0 ka (Fig. 1; Tables S1 and
S2 in the Supplement). Punctuating early Holocene deglaciation was deposition
of an extensive moraine system during a period locally referred to as the
Kapisigdlit stade. Although no direct moraine ages exist, deposition of the
Kapisigdlit stade moraines likely occurred sometime between ∼10.4 and
10.0 ka based on the timing of deglaciation from regional
radiocarbon and 10Be constraints and a single maximum-limiting
radiocarbon age of 10.17±0.34calka from reworked marine
sediments (Weidick et al., 2012; Larsen et al., 2014; Table S1). Following
early Holocene deglaciation, retreat of the GrIS continued inboard of its
current margin before readvancing during the late Holocene. In the study
area, the GrIS reached its historical maximum extent during the early to
mid-18th century (Weidick et al. 2012), which is marked by a prominent moraine
and trimline (Weidick et al., 2012; Figs. 1 and 2). In some locations the
GrIS still resides at or near its historical maximum extent (Kelley et al.,
2012), whereas in the KNS forefield, GrIS retreat from the historical maximum
is slightly more pronounced and has exposed fresh bedrock surfaces.
Field methods
Fieldwork was completed in 2017 CE and was primarily concentrated in the KNS
forefield and north of KNS at Qamanaarsuup Sermia (Figs. 1–3). In addition,
we collected samples for 10Be dating near Narsap Sermia, located
∼55km north of KNS, and near Kangaasarsuup Sermia located ∼20–25 km southwest of KNS (Lesnek et al., 2020; Fig. 1). Moraine
crests were mapped prior to fieldwork and updated in the field. This mapping
follows previous efforts (Weidick, 1974; Weidick et al., 2012; Pearce et al.,
2018), with the exception that we distinguish between early to middle Holocene
moraines and moraines marking the GrIS historical maximum extent (Figs. 2 and
3). Across the broader KNS region, the distinction is obvious. Moraines and
trimlines attributed to the historical maximum extent of the GrIS are close
to the modern ice margin and are fresh in appearance due to a lack of
vegetation and lichen cover. Early Holocene moraines are typically located
well outboard of the historical moraines and have extensive lichen cover. The key
exceptions to this spatial relationship are regions where ice is
topographically confined near small outlet glaciers (Figs. 1 and 3). In these
locations, the historical maximum and early Holocene moraines are closely
stacked yet are still easily distinguishable based on their morphologies and
degree of lichen cover (Figs. 3 and 4).
(a) View to the northeast in the
Qamanaarsuup Sermia region (Fig. 3). In the foreground is
a Kapisigdlit stade moraine crest resting directly adjacent to the
historical maximum extent of this sector of the GrIS (dashed line). The GrIS
is in the background, and there has been minimal retreat from the
historical maximum extent here. (b) View to the southwest in the
Qamanaarsuup Sermia region showing Marshall Lake and Goose Feather
Lake. Note the color contrast between the two lakes. Marshall Lake currently
does not receive meltwater from the GrIS. Goose Feather Lake is currently a
proglacial lake that receives silt-laden GrIS meltwater; the lake catchment
currently extends beneath the modern GrIS footprint (Figs. 3a and 6).
Samples for cosmogenic nuclide analysis were collected using a Hilti brand
AG500-A18 angle grinder–circular saw with diamond bit blades, as well as a hammer
and chisel. Sample locations and elevations were recorded with a handheld
GPS device with a vertical uncertainty of ±5m, and topographic
shielding was measured using a handheld clinometer. GPS units were
calibrated to a known elevation each day, either sea level or the stated
elevation of a lake derived from topographic maps.
Sediment cores from two proglacial–threshold lakes at Qamanaarsuup Sermia were
collected using a universal percussion corer and a Nesje-style
percussion–piston coring device (Fig. 3). Goose Feather Lake (informal name)
is located ∼2km from the GrIS margin and currently receives
GrIS meltwater. We collected two piston cores at a water depth of
12.60 m (17GOOF-A3 and 17GOOF-A4; 64.45328∘ N,
49.44373∘ W). Marshall Lake (informal name) is located ∼1km from the GrIS margin and does not presently receive meltwater
from the GrIS. We collected two cores from Marshall Lake with the universal
percussion corer system at a water depth of 5.95 m (17MAR-A2 and
17MAR-C1; 64.46361∘ N 49.44373∘ W).
10Be and 26Al geochemistry and accelerator mass spectrometric measurements
We completed 61 10Be and 12 26Al measurements;
54 of the 10Be samples and all of the 26Al samples were
processed at the Lamont–Doherty Earth Observatory (LDEO) cosmogenic dating
laboratory (Tables S3, S4, and S5 in the Supplement). The remaining seven
10Be samples were processed at the University at Buffalo Cosmogenic
Isotope Laboratory (Tables S3 and S4). In both laboratories, quartz separation as well as Be and Al isolation followed well-established protocols (Schaefer et al.,
2009). We quantified the amount of native 27Al in each quartz aliquot and
then added varying amounts of 27Al carrier to ensure that ∼1400–1750 mg of 27Al was achieved (Table S5). Total 27Al was
quantified after sample digestion using inductively coupled plasma optical
emission spectrometry analysis of replicate aliquots. Accelerator mass spectrometric analysis for
10Be samples was split between the Purdue Rare Isotope Measurement
(PRIME) Laboratory (n=36) and the Center for Accelerator Mass Spectrometry
at Lawrence Livermore National Laboratory (LLNL-CAMS; n=25); all
26Al samples were measured at PRIME.
All 10Be samples were measured relative to the 07KNSTD standard with a
10Be/9Be ratio of 2.85×10-12 (Nishiizumi et al., 2007),
and 26Al samples were measured relative to the KNSTD standard with
the value of 1.82×10-12 (Nishiizumi, 2004). For 10Be
samples measured at PRIME, the 1σ analytical error ranged from
1.9 % to 3.9 % with an average of 2.7%±0.5 % (n=36;
Table S3), and the 1σ analytical error for 26Al measurements
ranged from 1.9 % to 5.0 % with an average of 3.8%±0.9 %
(n=12; Table S5). For 10Be samples measured at LLNL-CAMS,
1σ analytical error ranged from 1.6 % to 4.2 %, with an average
of 2.3%±0.8 % (n=25; Table S3). Process blank corrections for
all 10Be and 26Al samples were applied by taking the
batch-specific blank value (expressed as the number of atoms) and subtracting this
value from the sample atom count (Tables S4 and S5). In addition, we propagate
through a 1.5 % uncertainty in the carrier concentration when calculating
10Be concentrations. We assume half-lives of 1.387 and
0.705 Ma for 10Be and 26Al (Chmeleff et al., 2010;
Nishiizumi, 2004).
In situ 14C measurements
We completed 12 in situ 14C extractions at the LDEO cosmogenic
dating laboratory following well-established LDEO extraction procedures (Lamp
et al. 2019; Table S6). All measured fraction modern values
are converted to 14C concentrations following Hippe and Lifton
(2014). The LDEO in situ 14C extraction laboratory has
historically converted samples to graphite prior to measurement by accelerator mass spectrometry and
LLNL-CAMS; however, we have recently transitioned to gas-source measurements
with the AixMICADAS instrument at CEREGE, which can directly measure ∼10–100 µg C and largely removes the need for the addition of a
carrier gas added to typical in situ 14C samples (Bard et al.,
2015; Tuna et al., 2018). Here, two samples underwent traditional
graphitization and were measured at LLNL-CAMS (Table S6), whereas the
remaining 10 samples were measured with the CEREGE AixMICADAS instrument
(Table S6). Both sets of samples underwent the same in situ 14C
extraction and 14C sample gas clean-up procedures, with only the
samples measured at LLNL-CAMS undergoing an additional graphitization
procedure (Lamp et al., 2019). Because we use two different measurement
approaches and our extraction efforts span the transition between sample
graphitization and gas-source measurements, we briefly discuss our data
reduction methods for both sets of measurements (Table S6).
Samples 17GRO-14 and 17GRO-74 were measured at LLNL-CAMS with 1σ
analytical uncertainties of 2.2 % and 3.0 % (Table S6). In situ
14C concentrations were blank-corrected using a long-term mean blank
value of 116894±3730714C atoms with the uncertainty in
the blank correction propagated in quadrature (n=27; updated from Lamp
et al., 2019). In addition, we propagate an additional 3.6 % uncertainty
in 14C concentrations based on the long-term scatter in internal
graphite-based CRONUS-A standard measurements (698109±25380atomsg-1; n=13; updated from Lamp et al., 2019);
stated in situ 14C concentrations for samples measured at
LLNL-CAMS have total uncertainties of 7.7 % and 10.4 %, respectively
(Table S6).
Samples measured at CEREGE have 1σ analytical uncertainties that range
between 1.0 % and 2.8 % with a mean of 1.4%±0.5 %
(n=10; Table S6). For our gas-source measurements presented here and for
future lab measurements, we recharacterized our extraction and measurement
procedure with a new set of process blank and CRONUS-A standard measurements
(Table S6). We completed six process blank gas-source measurements at CEREGE
with values ranging from ∼73000–175 000 14C atoms
(Table S6). One blank measurement is anomalously high (174813±3582 atoms; Table S6) and we suspect this blank was contaminated by
the atmosphere during collection in a break-seal. The remaining blank values have a
mean of 85768±1207014C atoms (n=5), and we tentatively
suggest that removing the graphitization procedure may also remove a source of
14C that was contributing to LDEO background 14C blank
values. Here, we use running-mean blank values of 81094±6972 (n=4)
and 85768±1207014C atoms (n=5) to correct sample
14C concentrations and uncertainties in the blank corrections are
propagated in quadrature (Table S6). In addition, we made five CRONUS-A standard
measurements at CEREGE. Our CRONUS-A measurements are remarkably consistent,
with a mean value of 662132±9849atomsg-1 (n=5;
1.5 % uncertainty), and are comparable to our graphite-based value of 698109±25380atomsg-1 (updated from Lamp et al.,
2019). Nonetheless, despite the promising consistency of our gas-source
CRONUS-A measurements, we conservatively propagate through an additional
3.6 % uncertainty in our sample concentrations based on the scatter in
long-term LDEO CRONUS-A measurements. Total 14C concentration
uncertainties for samples measured at CEREGE range from 4.3 % to 5.2 %
with a mean uncertainty of 4.6%±0.2 % (Table S6).
10Be and in situ 14C age calculations
10Be and in situ 14C surface exposure ages are
calculated using the Baffin Bay 10Be production-rate calibration
dataset (Young et al., 2013a) and the West Greenland in situ 14C
production-rate calibration dataset (Young et al., 2014). All ages are
presented using time-variant “Lm” scaling (Lal, 1991; Stone, 2000), which
accounts for changes in the magnetic field, although these changes are minimal
at this high latitude (∼64∘ N); using “St”
scaling, which does not account for changes in the magnetic field, results in
almost identical ages (<10 years) because the calibration sites are all
located at high latitudes. The 10Be and in situ 14C
calibration datasets are both derived from sites in western Greenland with
early Holocene exposure histories, and the in situ 14C calibration
measurements are derived from the same geologic samples as one of the
10Be calibration sites (Young et al., 2014). This combination of
calibration datasets ensures that the production rates and the
14C/10Be production ratio are regionally constrained. All ages
are calculated in MATLAB using code from version 3 of the exposure
age calculator found at https://hess.ess.washington.edu/ (last access: 27 August 2020), which implements an updated treatment of
muon-based nuclide production (Balco et al., 2008; Balco, 2017). We do not
correct nuclide concentrations for snow cover or subaerial surface erosion;
samples are almost exclusively from windswept locations, and many surfaces
still retain primary glacial features. In addition, we make no correction for
the potential effects of isostatic rebound on nuclide production because both
the production-rate calibration sites and sites of unknown age have
experienced similar exposure and uplift histories (i.e., the correction is
“built in”; Young et al., 2020a, b). Individual 10Be and in
situ 14C ages are presented and discussed with 1σ
analytical uncertainties, and moraine ages exclude the 10Be
production-rate uncertainty when we compare them to other 10Be-dated
features. When moraine ages are compared to independent records of climate
variability or ice-sheet change, the production-rate uncertainty is propagated through in quadrature (1.8 %;
Young et al., 2013a). To allow for direct
comparison to traditional radiocarbon constraints in the region, all
10Be and in situ 14C surface exposure ages are presented
in thousands of years BP (1950 CE); exposure ages relative to the year of
sample collection can be found in the Supplement (2017 CE; Tables S3 and S6).
Traditional 14C ages from proglacial–threshold lakes
Five radiocarbon ages from aquatic macrofossils were obtained from Marshall
Lake (Figs. 3 and 4; Table S7 in the Supplement), and one radiocarbon age from
an aquatic macrofossil was obtained from lake Kap01 (Fig. 1; Table S7). In
addition, we discuss two previously reported radiocarbon ages from Goose
Feather Lake, located adjacent to Marshall Lake (Lesnek et al., 2020;
Table S7). Aquatic macrofossils were isolated from surrounding sediment using
deionized water washes through sieves. Samples were freeze-dried and sent to
the National Ocean Sciences Accelerator Mass Spectrometry Facility (NOSAMS) at
Woods Hole Oceanographic Institution for age determinations. We targeted
aquatic macrofossils for dating because terrestrial macrofossils may persist
on the relatively low-energy Arctic landscape for hundreds of years before
washing into a lake basin; dating terrestrial macrofossils could skew our
interpretations. Hard-water effects on the 14C ages, which could
make age determinations erroneously old, are unlikely in our study area
because lake catchments are dominated by Archean gneiss, and the study lakes
are all well above local marine limits. All new radiocarbon ages are
calibrated using CALIB 8.2 and the INTCAL20 dataset, and previously
reported radiocarbon ages are recalibrated in the same manner using the
INTCAL20 and MARINE20 datasets (Stuiver et al., 2020; Reimer et al., 2020;
Heaton et al., 2020; Tables S1 and S7).
Ice-sheet model simulations of southwestern GrIS
change
We utilize recent paleo-simulations of southwestern GrIS change using the
high-resolution Ice Sheet and Sea-level System Model (ISSM; Larour et al.,
2012; Cuzzone et al., 2018, 2019; Briner et al., 2020). The model setup has
been previously described in Cuzzone et al. (2019) and Briner et al. (2020),
but here we briefly describe model attributes. The model domain extends from
the present-day coastline to the GrIS divide. The northern and southern
boundaries of the domain are far to the north and south of our study area.
The model resolution relies on anisotropic mesh adaptation to produce an
unstructured mesh that varies based on bedrock topography; bedrock topography
is from BedMachine v3 (Morlighem et al., 2017). For the southwestern GrIS,
high horizontal model mesh resolution is necessary in areas of complex bed
topography to prevent artificial ice-margin variability resulting from
interaction with bedrock artifacts that occur at coarser resolution (Cuzzone
et al., 2019). Thus, the model mesh varies from 20 km in areas where
gradients in the bedrock topography are smooth to 2 km in areas where
bedrock relief is high. In the KNS region the mesh varies from 2 to
8 km.
The ice model applies a higher-order approximation (Blatter 1995; Pattyn 2003)
to solve the momentum balance equations and an enthalpy formulation (Ashwanden
et al., 2012), with geothermal heat flux from Shapiro and Ritzwoller (2004),
to simulate the thermal evolution of the ice. Quadratic finite elements
(P1 × P2) are used along the z axis for the vertical interpolation,
which allows the ice-sheet model to capture sharp thermal gradients near the
bed, while reducing computational costs associated with running a linear
vertical interpolation with increased vertical layers (Cuzzone et al.,
2018). Sub-element grounding-line migration (Serrousi et al., 2013) is included
in these simulations; however, due to prohibitive costs associated with
running a higher-order ice model over paleoclimate timescales these
simulations do not include calving parameterizations or any submarine melting
of floating ice.
Nine ice-sheet simulations are forced with paleoclimate reconstructions from
Badgeley et al. (2020), who used paleoclimate data assimilation to merge
information from paleoclimate proxies and global climate models. The
temperature reconstructions rely on oxygen isotope records from eight ice
cores; the precipitation reconstructions use accumulation records from five
ice cores, and all are guided by spatial relationships derived from the
transient climate model simulation TraCE-21ka (Liu et al., 2009; He et al.,
2013). The climate reconstructions are shown to be in good agreement with
independent paleoclimate proxy data (Badgeley et al., 2020, and references
therein). Along with a main temperature and precipitation reconstruction,
Badgeley et al. (2020) provide two sensitivity precipitation reconstructions
due to uncertainty in the accumulation records and four sensitivity
temperature reconstructions due to uncertainty in the relationship between
oxygen isotopes and surface air temperature. Briner et al. (2020) pair three
of the temperature reconstructions with each of the three precipitation
reconstructions to yield nine combinations that are used as transient climate
boundary conditions to force nine ice-sheet simulations. Two of the five
temperature reconstructions were not used because they yield Younger Dryas
ice-sheet margins that are inconsistent with geologic data.
We use a positive degree day (PDD) method (Tarasov and Peltier, 1999) to
compute the surface mass balance from temperature and precipitation, and we use
degree day factors of 4.3 mm∘C-1d-1 for snow and
8.3 mm∘C-1d-1 for ice, with allocation for the
formation of superimposed ice (Janssens and Huybrechts, 2000). We use a lapse
rate of 6 ∘Ckm-1 to adjust the temperature of the climate
forcings to the ice-surface elevation.
Results
Adjacent to Qamanaarsuup Sermia, 27 10Be ages from moraine boulders
and boulders perched on bedrock range from 20.34±0.45ka to
8.91±0.20ka (Figs. 3 and 5; Table S3). Sediments in Goose
Feather Lake are composed of a lower gray silt unit overlain by organic
sediments, which is in turn overlain by gray silt (Lesnek et al., 2020). A
single radiocarbon age from bulk sediments at the basal sediment contact is
8280±90calyrBP, and a radiocarbon age from aquatic
macrofossils at the upper contact between organic and minerogenic sediments is
820±80calyrBP. Sediments in Marshall Lake display the same
silt–organic–silt stratigraphy as sediments in Goose Feather Lake (Fig. 6). A
radiocarbon age from aquatic macrofossils from the basal sediment contact is
8720±350calyrBP, and a radiocarbon age from aquatic
macrofossils at the uppermost contact is 520±20calyrBP
(Fig. 6). Three additional radiocarbon ages from aquatic macrofossils between
the lowermost and uppermost contacts are 7250±70, 3650±50, and 940±20calyrBP and are in
stratigraphic order (Fig. 6).
Representative boulder samples from the
Qamanaarsuup Sermia region (Fig. 3). Samples 17GRO-16,
17GRO-32, and 17GRO-39 are moraine boulder samples. 17GRO-25 is an erratic boulder
perched on bedrock inside the Kapisigdlit stade moraine located
only a few meters outboard of the historical maximum moraine, which can be
seen in the background. Samples 17GRO-46 and 17GRO-47 are erratic
boulders resting on bedrock located outboard of the Kapisigdlit
stade moraine.
(a) Sediment cores and calibrated radiocarbon ages
(±2σ) from Marshall (MAR) and Goose Feather (GOOF)
lakes. Note the distinct color differences between silt and organic
sediments (see Figs. 3 and 4). Details for radiocarbon ages can be found
in Table S7. (b) Sub-ice topography in the Qamanaarsuup
Sermia region generated using the BedMachine v3 digital elevation model (Morlighem et al., 2017)
compared to our chronology of ice-margin change developed from
10Be ages and radiocarbon-dated lake sediments (panel
a). Shown is the average 10Be age from each feature
on the landscape (including production-rate uncertainty; see Fig. 3) and
basal radiocarbon ages from panel (a). Deglaciation of the
landscape just outboard of the Kapisigdlit stade moraines occurred
at 10.29±0.23ka, followed by moraine deposition at 9.57±0.38ka. Ice retreated behind the modern margin at 9.29±0.18ka, but the ice margin remained within the drainage catchment of Goose
Feather Lake until 8.28±0.09calka before retreating farther
inland. The dashed line delimits the topographic threshold (T) under the
modern GrIS that the ice margin must cross in order for Goose Feather Lake
to receive silt-laden meltwater. Inflow of meltwater ceases when the GrIS
margin retreats behind this topographic threshold, which rests
∼1km behind the modern margin.
In the KNS region, 27 10Be ages from moraine boulders, erratics
perched on bedrock, and abraded bedrock surfaces range from 23.93±0.53 to 5.38±0.23ka (Table S3), and 12 in
situ 14C ages from bedrock range from 10.11±0.89 to 5.62±0.84ka (Table S6). In
addition, 12 26Al–10Be ratios range from 7.35±0.33
to 6.01±0.25 (all bedrock). A single radiocarbon age from aquatic
macrofossils at the basal contact between silt and organic sediments in lake
Kap01 is 9450±440calyrBP (Fig. 1; Table S7).
North of Narsap Sermia near Caribou Lake (Fig. 1), three 10Be ages
from boulders perched on bedrock are 9.07±0.32, 8.66±0.31, and 8.66±0.31ka (Fig. 1;
Table S3). South of KNS at Deception Lake, two 10Be ages from
boulders perched on bedrock are 10.66±0.34 and 9.52±0.32ka, and near One-way lake, two 10Be ages from
boulders perched on bedrock are 16.43±0.49 and 7.72±0.26ka (Fig. 1; Table S3).
Deglaciation chronologiesQamanaarsuup Sermia
A total of 27 10Be ages from the Qamanaarsuup Sermia region range
20.34±0.45 to 8.91±0.20ka (Figs. 3 and
5; Table S3); however, the 10Be ages are from three distinct
morphostratigraphic units. Here, the Kapisigdlit stade is marked by numerous
closely spaced moraine crests located immediately outboard of the historical
moraines. Three 10Be ages from boulders resting on bedrock located
outside the entire Kapisigdlit moraine suite are 10.39±0.32, 10.34±0.31, and 10.12±0.29ka, and they have a mean age of 10.29±0.14ka,
which serves as a maximum-limiting age for the Kapisigdlit stade moraines in
the region. Three 10Be ages from boulders resting on bedrock
immediately inboard of all Kapisigdlit stade moraines, but outboard of the
historical moraine, are 9.35±0.27, 9.31±0.30, and 9.21±0.26ka; they provide a
minimum-limiting age of 9.29±0.07ka (Fig. 3). Of the 21
10Be ages from Kapisigdlit stade moraine boulders, 5
10Be ages are likely influenced by 10Be inheritance as
they are older than the maximum-limiting 10Be ages and similar in
age to deglacial constraints found ∼140km west at the modern
coastline (11.67±0.32 and 11.01±0.41ka; Figs. 1 and 3; Table S3), or they date to when the GrIS
margin was likely situated >140km to the west somewhere on the
continental shelf (20.34±0.45, 16.72±0.40, and 14.63±0.36ka; Figs. 1 and 3;
Table S3). The remaining 16 10Be ages from the Kapisigdlit moraine
set show no trend with distance from the ice margin. These 10Be ages
overlap at 1σ uncertainties with each other, the
minimum-limiting 10Be ages, or the maximum-limiting 10Be
ages, suggesting that deposition of this suite of moraine crests happened
within dating resolution (i.e., we cannot resolve the ages of different
moraine crests). Combined, the 16 10Be ages from moraine boulders,
excluding outliers, have a mean age of 9.57±0.33ka, which
is morphostratigraphically consistent with bracketing maximum- and
minimum-limiting 10Be ages of 10.29±0.14 and
9.29±0.07ka, respectively (Fig. 3). Including the
uncertainty in the 10Be production-rate calibration, 10Be
ages from the Qamanaarsuup Sermia region reveal that the GrIS margin
approached its modern extent at 10.29±0.23ka, deposited the
Kapisigdlit stade moraines at 9.57±0.38ka, and retreated
behind the position of the historical maximum at 9.29±0.18ka.
Alternating silt–organic–silt sediment packages found in Goose Feather and
Marshall lakes are typical of those found in proglacial–threshold lakes in
southwestern Greenland (Fig. 6; Briner et al., 2010; Larsen et al., 2015;
Young and Briner, 2015; Lesnek et al., 2020). Silt deposition occurs when the
GrIS margin resides within the lake catchment but does not override the lake,
feeding silt-laden meltwater into the lake. Organic sedimentation occurs when
the GrIS margin is not within the lake catchment and meltwater is diverted
elsewhere. Despite Goose Feather and Marshall lakes residing adjacent to each
other on the landscape, their radiocarbon ages suggest slightly different
ice-margin histories. Marshall Lake has a small and highly localized drainage
catchment and does not receive GrIS meltwater at present. In contrast, because
Goose Feather Lake currently receives GrIS meltwater, its drainage catchment
extends somewhere beneath the modern GrIS. Goose Feather Lake is fed by
meltwater sourced from an outlet glacier resting in an overdeepening, and once
ice thins below the valley edge, GrIS meltwater is diverted elsewhere, likely
indicating that the Goose Feather drainage divide resides near the modern ice
margin (e.g., Young and Briner, 2015; Lesnek et al., 2020). Indeed, sub-ice
topography in the Qamanaarsuup Sermia region reveals that the topographic
threshold dictating whether meltwater is diverted to Goose Feather Lake or
elsewhere is located within ∼1km of the modern ice margin
(Fig. 6). Despite the GrIS margin retreating behind the position of the
historical maximum position at 9.29±0.18ka, silt deposition
in Goose Feather Lake until 8280±90calyrBP indicates that
the ice margin remained within ∼1km of its present position
between ∼9.3 and 8.3 ka (Fig. 6).
Kangiata Nunaata Sermia
A total of 15 10Be ages in the KNS region constrain the timing of
deposition of the Kapisigdlit stade moraines and the timing of when the GrIS
retreated behind the eventual historical maximum limit (Figs. 2 and
7–9). West of the KNS terminus, three 10Be ages from erratic
boulders perched on bedrock located immediately outboard of the Kapisigdlit
stade moraine are 11.44±0.34, 11.15±0.34, and 10.23±0.42ka. Seven
10Be ages from moraine boulders range from 11.22±0.35 to 5.38±0.23ka, and two 10Be
ages from erratic boulders perched on bedrock located immediately inside the
moraine are 10.27±0.29 and 10.25±0.31ka (Fig. 7). Deglaciation of the outer coast at Nuuk
occurred at ∼11.2–10.7 ka, suggesting that the single
10Be age of 10.23±0.42ka from just outboard of
the Kapisigdlit stade moraine is likely the closest maximum-limiting age on
moraine deposition, and 10Be ages of 11.44±0.34 and 11.15±0.34ka are likely influenced by a slight amount
of inheritance (Figs. 1 and 7). This maximum-limiting 10Be age is
consistent with a maximum-limiting radiocarbon age of 10 170±340calyrBP from a bivalve reworked into Kapisigdlit stade till
located down-fjord (Weidick et al., 2012; Larsen et al., 2014; Fig. 1;
Table S1). The 10.23 ka age is also consistent with the
10Be age of boulders outboard of the Kapisigdlit stade moraines in
the Qamanaarsuup Sermia study area of ∼10.29ka. There is
significant scatter in our moraine boulder 10Be ages; the oldest
10Be age of 11.22±0.35ka is likely influenced by
isotopic inheritance, whereas the younger outliers of 6.73±0.18 and 5.38±0.23ka reflect
post-depositional boulder exhumation (Fig. 7). The remaining 10Be
ages from the Kapisigdlit stade moraine have a mean age of 10.24±0.31ka (n=4), which is supported by our minimum-limiting
10Be ages of 10.27±0.29 and 10.25±0.31ka from immediately inside the moraine (Fig. 7). Indeed,
10Be ages from erratic boulders perched on bedrock located
immediately inside moraines across southwestern Greenland typically provide
constraints that are nearly identical to tightly clustered 10Be ages
from moraine boulders (e.g., Young et al., 2011a, 2013b; Lesnek
and Briner, 2018). Furthermore, our statistically identical 10Be
ages from outboard and inboard of the Kapisigdlit stade moraine, as well as
from moraine boulders themselves, indicate that moraine deposition occurred
rapidly within the resolution of our chronometer. Including
the production-rate uncertainty, we directly date the Kapisigdlit stade
moraine to 10.24±0.36ka, and all available supporting
10Be and 14C ages further constrain moraine deposition to
∼10.4–10.2 ka.
Representative boulder samples related to the
Kapisigdlit stade moraine west of KNS. 17GRO-64 and 17GRO-66 are moraine
boulders, whereas 17GRO-67 and 17GRO-70 are erratic boulders resting on bedrock
located immediately inboard and outboard of the Kapisigdlit stade moraine.
On the east side of KNS, three 10Be ages from immediately outboard of
the historical maximum are 12.86±0.57, 10.28±0.24, and 10.00±0.24ka (Fig. 9). The
10Be age of 12.86±0.57ka is, again, almost
certainly influenced by inheritance as this 10Be age predates the
timing of deglaciation at the outer coastline. The remaining 10Be
ages of 10.28±0.24 and 10.00±0.24ka
are consistent with the 10Be ages from inside the Kapisigdlit stade
moraine (and outboard of the historical maximum limit) on the west side of the
fjord. Our new and previously published age constraints reveal that deposition
of the Kapisigdlit stade moraine in the KNS forefield occurred at
ca. 10.4–10.2 ka, followed by retreat of the GrIS within the
historical maximum limit shortly thereafter (Fig. 2). Any possible moraine
correlative with the ∼9.6ka moraine found at Qamanaarsuup
Sermia would have been overrun by the historical advance of KNS.
Lastly, a basal minimum-limiting radiocarbon age of 9450±440calyrBP from Kap01 is, within uncertainties, identical to a
previously reported basal radiocarbon age of 9850±290calyrBP from the same lake (Tables S1 and S7; Larsen et al.,
2014). We note that our new radiocarbon age is from aquatic macrofossils,
whereas the previously published radiocarbon age is from bulk sediments (humic
acid extracts). Despite the risk of bulk sediments yielding radiocarbon ages
that are too old, basal radiocarbon ages from Kap01 suggest that the offset
between macrofossil- and bulk-sediment-based radiocarbon ages is likely
minimal during the initial onset of organic sedimentation following landscape
deglaciation. We do not advocate for the use of bulk sediments to develop
down-core chronologies, but paired macrofossil–bulk sediment measurements from
the same horizon often yield similar or indistinguishable radiocarbon ages in
southwestern Greenland (e.g., Kaplan et al., 2002; Young and Briner, 2015),
suggesting that bulk sediments will not produce significantly erroneous
radiocarbon ages in this region. These similarities in southwestern Greenland
likely result from several factors: (1) a large fraction of humic acid
extracts are aquatic in origin (Wolfe et al., 2004); (2) southwestern
Greenland is composed almost entirely of crystalline bedrock, thereby
minimizing potential hard-water effects; and (3) there is no significant
accumulated carbon pool during the initial phase of ecosystem development
(i.e., Wolfe et al., 2004). This latter point may be particularly influential
in southwestern Greenland because this region rests well inboard of the GrIS
margin during glacial maxima (located on the continental shelf), resulting in
a landscape that is likely ice-covered for a significant fraction of each
glacial cycle. Furthermore, this sector of the GrIS is primarily warm-based
and erosive, thereby further minimizing the likelihood of old carbon
accumulating on the landscape at lower elevations.
Auxiliary sites
At our site near Narsap Sermia, located ∼55km north of KNS,
three 10Be ages from boulders perched on bedrock located outboard of
the GrIS historical maximum and inboard of the Kapisigdlit stade limit have a
mean age of 8.80±0.24ka (8.80±0.29ka
including the production-rate uncertainty; Table S3), consistent with a
minimum-limiting basal radiocarbon age of 7460±110calyrBP
(Lesnek et al., 2020; Fig. 1; Table S1). Near Kangaasarsuup Sermia, located
∼35km south of KNS, two 10Be ages from boulders
perched on bedrock outboard of the historical limit and inboard of the
Kapisigdlit stade limit are 10.66±0.34 and 9.52±0.32ka. An existing 10Be and traditional 14C
age of 9.95±0.19ka (n=2) and 8790±190calyrBP from locations slightly more distal from the ice
sheet suggest that our age of 10.66±0.34ka is perhaps
influenced by a small amount of isotopic inheritance (Larsen et al., 2014;
Fig. 1; Tables S2 and S3). The remaining age of 9.52±0.32ka
is consistent with a minimum-limiting basal radiocarbon age of 9210±190calyrBP from Deception Lake (Fig. 1; Lesnek et al., 2020).
Lastly, south of Kangaasarsuup Sermia near One-way Lake, two 10Be
ages from boulders perched on bedrock outboard of the historical limit are
16.43±0.49 and 7.72±0.26ka (Fig. 1;
Table S3). The older of these two 10Be ages is influenced by
isotopic inheritance, leaving a single 10Be age of 7.72±0.26ka as the only estimate for the timing of local
deglaciation.
10Be–14C–26Al measurements from the KNS forefield
Prior to interpreting triple 10Be–14C–26Al
measurements in abraded bedrock located between the historical maximum extent
and the current ice margin, we use our new chronology of early Holocene
ice-margin change and historical observations to quantify the maximum duration
of Holocene exposure our bedrock samples sites could have experienced. First,
we use 10Be ages from immediately outboard of the historical limit
on the northeastern side of KNS and 10Be ages from just inboard of
the Kapisigdlit stade moraine on the southwestern side of KNS to define the
potential onset of Holocene exposure at our
10Be–14C–26Al bedrock sites. 10Be ages
from outboard of the historical limit overlap at 1σ uncertainties
(n=4; excluding one outlier), and we calculate a mean age of 10.20±0.14ka (10.20±0.23ka with production-rate
uncertainty) as the earliest onset of exposure at our inboard bedrock sites
(Fig. 2). This age represents the timing of deglaciation immediately outboard
of the historical limit and, assuming the continued retreat of the GrIS
margin, the initial timing of exposure for the inboard bedrock sites (e.g.,
Young et al., 2016). Next, we capitalize on historical observations in the KNS
region that constrain ice-margin change beginning in the 18th century (Weidick
et al., 2012; Lea et al., 2014, and references therein). Based on scattered
first-person observations, the advance towards the eventual historical maximum
extent likely began by 1723–1729 CE and culminated in ca. 1750 CE, with initial
ice-margin thinning taking place at ca. 1750–1800 CE (Weidick et al.,
2012). Broadly supporting this record of ice-margin migration is an early
photograph by Danish geologist Hinrich Rink dated to sometime in the 1850s
(Fig. 10). The photograph depicts the front of KNS as seen from the northwest
and clearly delineates an existing historical maximum trimline, indicating that the
local GrIS historical maximum was achieved and initial thinning from this
maximum began prior to 1850 CE (Fig. 10; Weidick et al., 2012). Additional
first-person descriptions indicate that the KNS ice margin was more extended
than today between ca. 1850 CE and at least 1948 CE, punctuated by the 1920 CE
stade, which marks a significant readvance of the ice margin (Fig. 2). Aerial
photographs reveal that the ice margin was only a few tens of meters east of our
bedrock sites on the western side of KNS at 1968 CE, suggesting site
deglaciation shortly beforehand (Weidick et al., 2012). Our eastern
ice-marginal bedrock sites were likely ice-free in ca. 2000 CE based on satellite
imagery.
(a) Photograph looking up-fjord towards KNS taken some
time in the 1850s by Danish geologist Hinrich Rink (Weidick et al., 2012).
Our northeastern KNS field site (Fig. 9) is located on the distal side of
Nunaatarsuk. (b) Close-up of Akullersuaq (A) and Nunaatarsuk (N) that
captures the trimline marking the historical maximum extent of KNS.
The photograph is housed in the archives of the National Museum in Copenhagen.
A digital copy was graciously provided by O. Bennike.
The majority of our bedrock sites are directly adjacent to the ice margin, and
we assume that pre-imagery historical observations of ice-margin change apply
to both sampling regions because any differences in ice-covered and ice-free
intervals between the two sampling sites are likely negligible for our
purposes. The available historical constraints indicate that our bedrock sites
became ice-covered in ca. 1725 CE (historical maximum advance phase), and our
sites on the western side of KNS likely became ice-free in ca. 1968 CE; ice-marginal sites on the east side of KNS became ice-free in ca. 2000 CE. These
observations indicate that the western bedrock sites experienced 243 years of
historical ice cover, whereas the eastern sites experienced 275 years of ice
cover. With the earliest possible onset of exposure occurring at 10.20±0.23ka as constrained by our 10Be ages, we assume that
the maximum duration of Holocene surface exposure at all of our sites is
10.0 kyr. We do note, however, that three of our bedrock sampling
sites on the northeastern side of KNS are at a higher elevation and lie closer
to the historical maximum limit than the sites adjacent to the modern ice
margin (Figs. 2 and 9). These high-elevation sites almost certainly
experienced shorter historical ice cover than our lower-elevation ice-marginal
sites, likely on the order of ∼150–200 years. But considering the
uncertainties in our chronology and analytical detection limits, we assume
they have the same maximum Holocene exposure duration of
10 kyr. Lastly, samples could inherit cosmogenic nuclides from an
earlier exposure (i.e., inheritance), and in the strictest sense, the most
recent period of exposure for our bedrock sites equates to only the last few
decades. Because of the well-constrained maximum possible exposure duration
provided by geologic constraints and historical observations, here isotopic
inheritance refers to exposure ages older than 10.20±0.23ka
(i.e., pre-Holocene exposure).
Apparent in situ 10Be and 14C surface exposure
ages
Apparent 10Be ages from abraded bedrock surfaces on the northeastern
side of KNS, listed from just inboard of the historical maximum limit towards
the modern ice margin, are 17.46±0.47, 17.15±0.44, 12.95±0.31, 9.96±0.24, 19.83±0.40, 23.93±0.53, 15.16±0.40,
and 6.83±0.16ka (Figs. 2, 9, 11, and 12; Table S3). On the
southwestern side of KNS adjacent to the modern ice margin, 10Be
ages from abraded bedrock surfaces are 6.94±0.16, 6.75±0.16, 6.56±0.27, and 6.51±0.14ka, all roughly at equal distance from
the present ice margin (Figs. 2 and 13; Table S3). Along our northeastern
transect, six of the eight apparent 10Be ages exceed the maximum allowable
Holocene exposure duration (10.20±0.23ka). Apparent
10Be ages greater than this indicate the presence of inherited
10Be accumulated from a period of pre-Holocene exposure and
insufficient subglacial erosion during the last glacial cycle to reset the
cosmogenic clock. Of the two remaining 10Be ages not influenced by
isotopic inheritance, the 10Be age of 9.96±0.24ka
is statistically identical to the maximum allowable duration of Holocene
exposure for the landscape located between the historical moraine and the
modern ice margin. In addition, a 10Be age of 6.83±0.16ka directly adjacent to the modern margin is suggestive of less
exposure and more burial (or more erosion; see Sect. 5.2) at this site, and it is
also statistically identical to apparent 10Be ages from the
southwestern side of KNS, indicating similar Holocene exposure histories.
Sampled bedrock surfaces located between the historical
maximum extent of the GrIS and the modern ice margin on the northeastern
side of KNS.
Sampled bedrock surfaces adjacent to the modern ice
margin on the northeastern side of KNS.
Although the maximum amount of allowable Holocene exposure is
well-constrained, we pair our 10Be measurements with in situ
14C measurements to (1) further assess the magnitude of isotopic
inheritance in our bedrock samples and (2) constrain post-10 ka
fluctuations of the GrIS margin. Whereas long-lived nuclides such as
10Be must be removed from the landscape via sufficient subglacial
erosion, the relatively short half-life of 14C (t1/2=5700 years) allows previously accumulated in situ 14C to
decay to undetectable levels after ∼30kyr of simple burial of a
surface by ice; with the aid of subglacial erosion, in situ 14C
can reach undetectable levels more quickly. In contrast to our apparent
10Be ages, which display varying degrees of Holocene exposure and
isotopic inheritance, in situ 14C measurements are consistent with
Holocene-only exposure histories. On the southwestern side of KNS, four
apparent in situ 14C ages range from 6.59±0.47 to 5.62±0.84ka, and paired
10Be–14C measurements yield concordant exposure ages
(Fig. 13). Within our northeastern bedrock transect, the two highest-elevation
samples near the historical maximum limit with 10Be ages of 17.46±0.47 and 17.15±0.44ka have significantly
younger in situ 14C ages of 10.11±0.89ka and 9.46±0.78ka, respectively (Figs. 2 and 9). The next sample along
this transect has a 10Be age of 12.95±0.31ka and
an in situ 14C age of 9.70±0.80ka, followed by
a sample with concordant 10Be and in situ 14C ages of
9.96±0.24ka and 9.80±0.98ka, respectively
(Figs. 2 and 9). Within our cluster of samples closest to the ice margin along
the northeastern transect, the one apparent 10Be age of 6.83±0.16ka is matched by a statistically identical in situ
14C age of 6.99±0.50ka, indicating that this
10Be age is likely not influenced by isotopic inheritance (Figs. 2
and 9). The remaining three samples in this cluster with 10Be ages
of 23.93±0.53, 19.83±0.40, and 15.16±0.40ka have significantly younger in situ 14C ages of
6.68±0.48, 6.29±0.73, and 6.49±0.46ka,
respectively (Figs. 2 and 9).
(a)10Be–14C apparent exposure
ages from the northeastern and southwestern sides of KNS plotted against
sample elevation. Apparent 10Be ages that are older
than ∼10.3ka are influenced by isotopic inheritance, yet
their corresponding in situ 14C
ages are younger and consistent with Holocene-only exposure histories.
In situ 14C ages vs. sample elevation are consistent with ice-margin thinning. (b) Paired
14C–10Be diagram. The x axis
is the measured 10Be concentration normalized by the
site-specific production rate (years); the y axis is the measured
14C/10Be ratio normalized to
the production ratio. We use a regionally constrained
14C/10Be spallation production
ratio of 3.12 (Young et al., 2014). The simple exposure region (black lines)
is defined by the continuous and steady-state erosion lines. Samples that
are influenced by isotopic inheritance are excluded because they yield
artificially low and meaningless ratios. Remaining samples are all
consistent with constant exposure; we estimate a minimum burial detection
limit of 625 years.
There are two modes of apparent in situ 14C exposure ages: a
cluster of in situ 14C ages at ∼10ka and a
second cluster at ∼6–7 ka (Figs. 2 and 14). Perhaps more
importantly, however, is that these two modes correlate with two distinct
morphostratigraphic surfaces. In situ 14C ages of ∼6–7 ka are all from sites located directly adjacent to the
modern ice margin. On the southwestern side of KNS, in situ 14C
ages of ∼6–7 ka are matched by concordant 10Be
ages (Figs. 2 and 14). On the northeastern side of KNS, in situ
14C ages of ∼6–7 ka are found closest to the
modern ice margin. One of these sites has concordant 10Be and in
situ 14C ages, while at the remaining sites, in situ
14C ages of 6–7 ka are significantly younger than their
paired 10Be ages, despite all sample sites appearing to have
undergone significant subglacial erosion (Figs. 2 and 14). In situ
14C ages of ∼10ka only exist at our
high-elevation sites directly adjacent to the historical maximum
limit. Moreover, in situ 14C ages that range between 10.11±0.89 and 9.46±0.78ka at our high-elevation
sites are statically indistinguishable from the maximum duration of Holocene
exposure these sites could have experienced (∼10kyr; Figs. 2
and 14). Thus, the in situ 14C ages of ∼10ka,
including the paired 10Be–14C ages of ∼10ka, indicate that following deglaciation of the landscape just
outboard of the historical moraine at 10.20±0.23ka, the GrIS
margin continued to retreat inland and expose our high-elevation bedrock sites
immediately thereafter. Moreover, subglacial erosion during the brief period
of historical ice cover was negligible at these high-elevation sites because
any significant subglacial erosion would result in apparent in situ
14C ages that are younger than the maximum Holocene exposure history
these sites could have experienced. Our younger 10Be and in situ
14C ages of 6–7 ka, on the other hand, reflect some
combination of less Holocene exposure and/or more subglacial erosion than the
high-elevation sites.
14C/10Be ratios, transient burial, and subglacial erosion
Combining two or more nuclides with different half-lives quantifies the
integrated amount of surface exposure and burial (e.g., Bierman et al., 1999;
Goehring et al., 2011). Following a period of surface exposure, burial of a
surface by overriding ice will cease nuclide production and lead to the faster
decay of the short-lived nuclide relative to the nuclide with a longer (or
more stable) half-life. Typically, the longer-lived nuclide is used to constrain
the total amount of surface exposure, whereas the nuclide with a shorter
half-life functions as the burial chronometer. Because the production ratio of
two nuclides of a constantly exposed surface is known, a measured sample ratio
below the constant production value represents the duration of surface
burial. Over the Holocene timescale considered here, 10Be functions
as an essentially stable nuclide because of its long half-life
(t1/2=1.387Myr), whereas 14C, with its short half-life
(t1/2=5700 years), acts as the burial chronometer (e.g., Goehring et al.,
2011).
We only consider measured 14C/10Be ratios in samples that do not
have inherited 10Be; inherited 10Be results in physically
unobtainable ratios. Using the average precision of our measured
14C/10Be ratios (5.7 %), we estimate a minimum burial
detection limit resolvable to 625 years at 1σ (Table S8 in the
Supplement). All six samples without 10Be inheritance have
14C/10Be ratios that are indistinguishable from constant exposure
(Fig. 14; Table S8). One of these pairings with concordant
14C–10Be ages of ∼10ka on the
northeastern side of KNS indicates that the high-elevation bedrock sites
likely became reoccupied by the GrIS only once as the ice sheet advanced
towards the historical maximum extent <625 years ago (Figs. 2 and 14). The
remaining samples with 14C–10Be ages of ∼6–7 ka also reveal <625 years of ice burial (Fig. 14). Our
measured 14C/10Be ratios are consistent with the observation of
∼245–275 years of recent historical ice cover as being the only period
of ice cover these sites experienced after early Holocene deglaciation. Our
measured ratios, however, cannot rule out brief pre-18th century advances of
the GrIS that may have covered our ice-marginal bedrock sites. Because our ice-marginal sites reside so close to the modern ice margin, it is likely that any
brief advance of KNS prior to the 18th century would have covered our sampled
bedrock locations.
Our measured 14C–10Be ratios do not indicate any
significant amounts of burial at our ice-marginal sites, but concordant
14C–10Be ages of ∼6–7 ka suggest an
exposure history and/or degree of subglacial erosion fundamentally different
than the high-elevation 10 ka landscape (Figs. 2 and 14). The
simplest interpretation is that the concordant 14C–10Be
ages and constant production ratios reflect one period of surface exposure
over the last ∼6–7 kyrBP, prior to the period of historical
ice cover. In this scenario, deglaciation of the high-elevation bedrock sites
occurred at 10.20±0.23ka, but instead of continued inland
retreat of the GrIS margin, the ice margin stabilized for several thousand
years and then retreated inboard of today's margin around
6–7 ka. However, another possibility is that these sites also
deglaciated at 10.20±0.23ka or later but later experienced
significant subglacial erosion during the period of historical ice
cover. Subglacial erosion through the production–depth profile in bedrock
would result in younger apparent exposure ages (e.g., Goehring et al., 2011;
Young et al., 2016) while maintaining a 14C/10Be ratio consistent
with constant exposure as long as erosion was limited to the first few
tens of centimeters over which spallation dominates nuclide production.
Modeled 10Be (black line)
and in situ 14C (green line)
production with depth in bedrock using a typical rock density of 2.65 gcm-3 for gneiss. A normalized
10Be and in situ 14C concentration of 1 equals a prescribed exposure
duration of 10, 9, 8, or 7.5 kyr. Symbols mark the depth below
the prescribed deglacial surface along the production–depth profile that the
GrIS would have to erode to during the period of historical ice cover that
results in our measured nuclide concentrations. The
10Be production–depth profile is not fit to the data
points and any 10Be measurement will fall somewhere
along the depth profile. Bars on the left (10Be) and
right (14C) sides of the diagram are the average
erosional depths (±1SD) of each exposure–erosion scenario (see
Table S9). For figure clarity, we only show where our individual
10Be measurements fall along the production–depth
profile; the average erosional depths based on the in
situ 14C measurements are on the right side
of the figure (Table S9). Production by spallation and muons are calculated
independently (muons according to Balco, 2017; model 1A). Z is the mass
depth below the surface (gcm-2) and is
the product of depth (cm) and material density (gcm-3). Pi is the production
rate of 10Be and in situ
14C (atomsg-1yr-1) by spallation or muons at depth z.
Pi (0) is the surface production rate via spallation
or muons. Λ is the effective attenuation length (gcm-2). Included here are only
10Be and in situ 14C concentrations for samples with concordant
apparent exposure ages of ∼6–7 ka (17GRO-71, 17GRO-72, 17GRO-73,
17GRO-74, and 17GRO-75; Tables S6 and S9).
To explore the possibility that concordant 10Be and 14C
ages of ∼6–7 ka are a product of significant subglacial
erosion, we cast apparent 10Be and 14C ages as a function
of total erosional depth into bedrock during the period of historical ice
cover assuming varying lengths of total surface exposure (Fig. 15; Table S9 in
the Supplement). Because of the significantly larger uncertainties in the in
situ 14C measurements, we only use the 14C-based erosion
depths as a check on the 10Be-based erosion depths. The more precise
10Be measurements result in more precise estimates of erosional
depth, and because we made paired 10Be–14C measurements
(versus single nuclide measurements in separate geological samples at
different locations), our 10Be measurements provide more robust
constraints of simulated erosional depth. As an estimate of the maximum
amount of total erosion our sample sites could have experienced, we assume
that our ice-marginal sites experienced 10 kyr of total exposure,
which is constrained by the timing of deglaciation from just outboard of the
historical moraines and the estimated duration of historical ice cover
(Sect. 5). We then project the 10Be and 14C
production–depth profiles in bedrock at each site and determine the depth
below the theoretical 10 kyr surface to which our measured
10Be and 14C concentrations equate (Fig. 15;
Table S9). Assuming a total exposure duration of 10 kyr,
10Be- and 14C-based total erosional depths are 24.8±1.8 and 20.1±4.1cm, respectively, during the
period of historical ice cover (Fig. 15; Table S9). Using the known duration
of historical ice cover (Sect. 5), these erosional depths translate to
abrasion rates of 1.00±0.11mmyr-1 (10Be) and 0.81±0.19mmyr-1
(14C). Using total exposure durations of 9, 8, and 7.5 kyr
yields 10Be-based erosion depths of 18.3±1.8, 11.1±1.8, and 7.2±1.8cm, respectively, which equate
to abrasion rates of 0.74±0.10, 0.45±0.08, and 0.29±0.08mmyr-1 (Fig. 15;
Table S9).
Calculated abrasion rates are all well above the canonical polar subglacial
abrasion rate of 0.01 mmyr-1 (Hallet et al., 1996) but notably
similar to abrasion rates inferred at the Jakobshavn Isbræ forefield
constrained by similar methodology (0.72±0.26mmyr-1; Young
et al., 2016); however, there are key differences between the two landscapes
that aid in further limiting the plausible abrasion rates in the KNS
forefield. Dozens of 10Be ages from the Jakobshavn Isbræ region,
most notably from bedrock inside and outside the Jakobshavn Isbræ
historical maximum limit, reveal a landscape entirely devoid of isotopic
inheritance suggestive of a highly erosive environment (Young et a., 2013b, 2016). In contrast, new 10Be measurements presented
here influenced by isotopic inheritance, including from bedrock, coupled with
previous 10Be measurements from the region (e.g., Larsen et al.,
2014), at least qualitatively point to a less erosive GrIS in the KNS
region. We think it is unlikely that subglacial abrasion rates at KNS can
match or exceed those from the Jakobshavn Isbræ forefield, and we therefore
favor an interpretation with less site exposure over significant
amounts of subglacial abrasion. Moreover, the in situ 14C ages
(and one 10Be age) from our high-elevation bedrock sites on the
northeastern side of KNS are indistinguishable from the timing of early
Holocene deglaciation and indicate that, at least at these high-elevation
sites, subglacial abrasion during historical ice cover was negligible.
The most straightforward process to generate statistically identical
10Be and in situ 14C ages across all of the ice-marginal
bedrock sites is for these sites to have undergone relatively little to no
subglacial abrasion following the period of early to middle Holocene
exposure. We also doubt that significant subglacial abrasion in the KNS
forefield during the period of historical ice cover would result in such
strikingly uniform apparent exposure ages (i.e., uniform abrasion rates). On
the other hand, it is unrealistic to assume that our ice-marginal sites did
not experience some degree of abrasion because this is not a cold-based ice
environment and striations were routinely observed at sampling locations;
these features were likely formed during the period of historical ice
cover. To estimate the likely timing of deglaciation at our ice-marginal sites
that have concordant 10Be and in situ 14C ages of
6–7 ka, we present a range of deglaciation estimates. To
constrain the latest possible age of deglaciation, we use all of the bedrock
10Be ages on the southwestern side of KNS (n=4), combined with the
apparent in situ 14C ages on the northeastern side of KNS that
have paired 10Be measurements influenced by inheritance (n=3) and
a single 10Be age that is not influenced by isotopic
inheritance. The mean age of these samples is 6.63±0.21ka,
which does not include the last ∼275 years of historical ice cover when
minimal to no (i.e., 14C) isotope production would have occurred,
and assumes zero erosion during historical ice cover. Accounting for the
period of historical ice cover, the timing of middle Holocene deglaciation
from our ice-marginal bedrock sites is 6.91±0.21ka. As an
upper bound on the timing of deglaciation, we rely on our modeled scenario of
7.5 kyr of exposure resulting in an abrasion rate of 0.29±0.08mmyr-1. As with our measured apparent 10Be and
in situ 14C ages, the modeled 7.5 kyr scenario does not
include the ∼275 years of historical ice cover. Including historical ice
cover results in a deglaciation age of 7.78 ka. Combined, our
favored interpretation is that the ice-marginal bedrock sites likely first
became ice-free sometime between ∼7.8 and ∼6.9ka and
experienced abrasion during the period of historical ice cover at no more than
∼0.3mmyr-1. We suggest that this estimate sufficiently
accounts for some degree of subglacial abrasion during the period of
historical ice cover as suggested by field observations, while also
acknowledging that tightly clustered apparent 10Be and in situ
14C ages are suggestive of bedrock surfaces that have undergone
minimal modification following initial mid-Holocene exposure.
26Al /10Be ratios
As with 14C/10Be ratios, 26Al/10Be ratios can be
used to measure integrated surface exposure and burial, albeit over much
longer timescales. The 26Al–10Be pairing utilizes
preferential decay of 26Al (t1/2=0.705Ma) relative
to 10Be (t1/2=1.387Ma) to quantify exposure and burial
over glacial–interglacial timescales or longer (e.g., Bierman et al., 1999;
Fabel et al., 2002; Gjermundsen et al., 2015). The canonical
26Al/10Be production ratio is considered to be 6.75 (Balco and Rovey,
2008; Balco et al., 2008), but measurements from western Greenland constrain
the 26Al/10Be production ratio to 7.3±0.3, suggesting that the
production ratio scales with latitude and elevation (Corbett et al.,
2017). Modeling suggests that the 26Al/10Be production ratio at
both sea level and high latitude is ∼7.0–7.1 (Argento et al., 2013).
Our 26Al/10Be ratios range from 7.39±0.33 to 6.01±0.25 in the KNS forefield (n=12; Table S5). Because of the extremely close
proximity of all of our bedrock samples, these surfaces must have the same
exposure and burial histories over glacial–interglacial timescales; the value
of 6.01±0.25 is anomalously low relative to our remaining
measurements. After removing the lowest ratio, remaining ratios range from
7.39±0.33 to 6.71±0.24 with a mean of 7.05±0.24 (n=11;
Table S5). Further limiting this dataset to samples that have no detectable
inherited 10Be results in a mean value of 7.17±0.20 (n=6;
Table S5). With the exception of our lowest measured ratio (6.01±0.25),
each of our 26Al/10Be ratios overlaps with the constant
production values at 2σ uncertainty regardless of which constant
production is used. Yet, our 26Al/10Be ratios are systematically
greater than 6.75 and suggest that the true production ratio is >6.75,
more consistent with recent modeled and empirical estimates (Table S5; Argento
et al., 2013; Corbett et al., 2017).
The burial detection limit for the paired 26Al–10Be method
is typically on the order of approximately one glacial cycle, although this is
dependent on the uncertainty in the measured 26Al/10Be ratio. All
of our measured 26Al/10Be ratios suggest constant exposure,
including samples with inherited 10Be (apparent 10Be ages
>10ka), yet the known ice-margin history requires that
constant exposure could have only occurred over the last ∼10kyr
(Fig. 16; excluding the brief period of historical ice cover that is
undetectable with the 26Al–10Be chronometer). The most
likely source of the excess 10Be in samples with apparent
10Be ages >10ka is surface exposure during Marine
Isotope Stage 5e (MIS; ∼129–116 ka; Stirling et al.,
1998). Brief exposure of our bedrock surfaces during MIS 5e and surface burial
between MIS 5d and ∼10ka, followed by re-exposure for the
last 10 kyr, is the most straightforward scenario to have measured
26Al/10Be ratios consistent with constant exposure while also
containing a slight amount of inheritance. Moreover, this scenario is
consistent with the broad outline of GrIS change over the last glacial
cycle. Greenland ice-core data and offshore sediment records reveal that the
GrIS was smaller than today during MIS 5e (Colville et al., 2011; NEEM, 2013),
which suggests that bedrock currently emerging from beneath the GrIS was
likely exposed for some period of time during MIS 5e.
Paired
26Al–10Be diagram with all
measurements from the KNS forefield (n=12). The x axis is the
measured 10Be concentration normalized by the
site-specific production rate (years); the y axis is the measured
26Al/10Be ratio normalized to
the production ratio. We use a production ratio of 7.07 (Argento et al.,
2013). The black lines define the simple exposure region. All of our
measured ratios, with the exception of 17GRO-13, indicate constant exposure
of the sample site (values listed in Table S5). Because of the close
proximity of these sites, they must have the same glacial–interglacial
exposure history, and thus we consider 17GRO-13 a likely outlier.
At the same time, we cannot rule out the possibility that small amounts of inherited
10Be, coupled with 26Al/10Be ratios consistent with
constant exposure, are a result of exposure during MIS 3. Terrestrial evidence
of a restricted GrIS during MIS 3 is limited so far to select sites in
northern Greenland (e.g., Larsen et al., 2018), and marine-based sediment
records from off southwestern Greenland point to significant MIS 3 recession
of the GrIS; however, the MIS 3 ice margin may have still been located off the
modern coastline, and thus our sample sites would remain ice-covered
(Seidenkrantz et al., 2019). In addition, any significant exposure of our KNS
bedrock sites during MIS 3 would require significant amounts of burial (or
erosion) in order to have any accumulated in situ 14C decay away
to undetectable levels, as there is no evidence of inherited in situ
14C in our samples. Moreover, eustatic sea level curves indicate
that sea level was at least 30–40 m below present during MIS 3
(Siddall et al., 2003; Grant et al., 2014). Most of this eustatic sea level
signature is driven by changes in the Laurentide Ice Sheet, but it is
difficult to imagine a complete decoupling of Laurentide and Greenland ice-sheet behavior whereby the Laurentide remains relatively large, while the GrIS is
smaller than today during MIS 3. It is certainly possible that our sites were
exposed during MIS 3, but we
favor the more straightforward explanation of MIS 5e exposure at our sample
site, considering the following: (1) ice-core records reveal that
the region was likely warmer during MIS 5e vs. MIS 3 (NGRIP, 2004; NEEM,
2013), (2) balancing the MIS 5e eustatic sea level budget likely requires a
significant contribution from Greenland (Dutton et al., 2015), and (3) there is a
lack of any additional terrestrial evidence in southwestern Greenland for an
MIS 3 ice-sheet configuration similar to or more restricted than today. While it is perhaps unsurprising that our 26Al–10Be
measurements in the KNS forefield suggest surface exposure during a previous
interglacial, these measurements nonetheless suggest the GrIS margin in the
KNS region was at or behind the present margin during MIS 5e.
Holocene evolution of the southwestern Greenland Ice SheetEarly Holocene moraine deposition
Direct 10Be ages from moraine boulders constrain Kapisigdlit stade
moraine deposition to 10.24±0.31ka, which is supported by
statistically identical bracketing 10Be ages from erratic boulders
located outboard and inboard of the Kapisigdlit stade moraine (Figs. 2 and
7). North of KNS at Qamanaarsuup Sermia, 10Be ages from moraine
boulders indicate that deposition of the so-called Kapisigdlit moraine occurred at
9.57±0.33ka, which is supported by bracketing
10Be ages of 10.29±0.14 and 9.29±0.07ka (Fig. 3). These two moraines share identical
maximum-limiting ages, but the direct moraine and minimum-limiting
10Be ages between the two sites are statistically
distinguishable. Maximum-limiting 10Be ages from erratic boulders
located immediately outboard of a moraine can provide a close constraint on
the age of moraine deposition if moraine deposition occurs via a brief
stillstand of the ice margin (e.g., Young et al., 2013b). If moraine
deposition occurs after a readvance of the ice margin, however, maximum-limiting 10Be ages need not be close limiting ages. In contrast,
minimum-limiting 10Be ages on erratic boulders from immediately
inside a moraine should provide close minimum-limiting constraints regardless
of whether moraine deposition occurred through a stillstand of the ice margin during
retreat or after a readvance. The distribution of 10Be ages
presented here suggests that deposition of the Kapisigdlit stade moraine in
the immediate KNS region occurred via a stillstand of the ice margin or a
brief readvance that occurred within the resolution of our chronometer. At
Qamanaarsuup Sermia, however, there is a gap of several hundred years between
maximum-limiting 10Be ages and moraine-based 10Be
ages, and a total of ∼1kyr between maximum- and
minimum-limiting 10Be ages suggests that the Qamanaarsuup Sermia
moraine suite was deposited after a readvance of the GrIS.
Another possibility is that the Qamanaarsuup Sermia moraine complex is an
amalgamation of moraines relating to chronologically distinct advances or
stillstands of the ice margin. The numerous and tightly packed moraines here,
vs. the more well-defined Kapisigdlit stade moraine at KNS, are suggestive of a
stagnating or oscillating ice margin. Moreover, our moraine boulder dataset
contains more scatter than we typically observe in southwest Greenland (e.g.,
Young et al., 2020a), suggesting it is possible that we sampled moraine
boulders from two or more distinct advances. In this case, combining all of our
10Be ages from moraine boulders at Qamanaarsuup Sermia would
inadvertently mask the timing of two or more advances; for example, if advances
occurred at ca. 10.4–10.3 and 9.3–9.0 ka, combining all
10Be ages might result in an average 10Be age of ∼9.7ka, especially if moraine boulders from each advance are
reworked by the ice margin. In several instances in southwest Greenland,
10Be ages from erratics just inboard of a moraine are statistically
identical to the moraine boulders themselves (e.g., Young et al., 2013b, 2020a), perhaps indicating that our minimum-limiting age of 9.29±0.07ka constrains an advance of the GrIS in this sector to ∼9.3ka (moraine closest to erratics), and moraines located farther
away from the ice margin might relate to an advance of the GrIS closer in age
to our maximum-limiting 10Be ages (10.29±0.14ka;
Fig. 3). Our 10Be ages from moraine boulders at Qamanaarsuup Sermia,
however, show no trend across moraines or with distance from the ice
margin. We prefer the more conservative interpretation acknowledging that
if two or more distinct advances occurred, we cannot resolve these advances
with our dataset. We can confidently say that all moraines were deposited
between 10.29±0.14 and 9.29±0.07ka,
and a moraine age of 9.57±0.33ka is consistent with these
bracketing ages.
Including the uncertainty in the 10Be production rate, the
Kapisigdlit stade moraine in the KNS forefield was deposited at 10.24±0.36ka, consistent with a maximum-limiting radiocarbon age of
10170±340calyrBP (Weidick et al., 2012; Larsen et al.,
2014), and moraine deposition at Qamanaarsuup Sermia occurred at 9.57±0.38ka. Note that these moraine ages overlap at 1σ only
when including the 10Be production-rate uncertainty, which results
in systematic shifts in age and is only needed when comparing these moraine
ages to independent chronometers; these moraines are distinguishable at
1σ in 10Be space. Moraine deposition in the KNS forefield at
10.24±0.36ka is, within resolution, synchronous with
widespread moraine deposition in southwestern Greenland and Baffin Island
at ca. 10.4–10.3 ka. Indeed, 10Be ages across several
locations in Baffin Bay reveal that sectors of the GrIS, Laurentide Ice Sheet,
and independent alpine glaciers on Greenland and Baffin Island all deposited
moraines at ca. 10.4–10.3 ka, likely in response to
freshwater-induced regional cooling (Young et al., 2020a). Contemporaneous
moraine deposition indicates that the KNS sector of the GrIS also likely
responded to regional cooling at ca. 10.4–10.3 ka. Moraine deposition
at 9.57±0.38ka at Qamanaarsuup Sermia, however, does not fit a
well-established regional pattern of moraine deposition. Similar to moraine
deposition at ca. 10.4–10.3 ka, widespread moraine deposition across
Baffin Bay occurred at ca. 9.3–9.0 ka and is thought to be driven by
the 9.3 ka cooling event displayed in Greenland ice cores (Young
et al., 2020a). The Qamanaarsuup Sermia moraine age (9.57±0.38ka) is, within uncertainties, synchronous with the
9.3 ka cooling event, but 10Be-dated moraines in Baffin Bay
consistently date at or slightly after the 9.3 ka cooling event
(Young et al., 2011b, 2020a; Crump et al., 2020). North of KNS,
10Be ages constrain deposition of a moraine to 9.7±0.7ka, consistent with moraine deposition at Qamanaarsuup Sermia
despite somewhat larger uncertainties (Lesnek and Briner, 2018). In addition
to widespread moraine deposition in Baffin Bay at ca. 10.4–10.3 and
9.3–9.0 ka, these emerging 10Be ages tentatively
suggest an additional mode of moraine deposition at ca. 9.7 ka, perhaps
in a response to freshwater-related cooling (Lesnek and Briner,
2018). Regardless of our ability to correlate the Qamanaarsuup Sermia moraine
dated to 9.57±0.38ka with moraines beyond the KNS region, our
10Be ages reveal that early Holocene moraines across the broader KNS
region are not equivalent features. Our results reveal at least two periods of
moraine deposition occurred at 10.24±0.36 and 9.57±0.38ka. Lastly, we note that there is no moraine associated with
the 8.2 ka abrupt cooling event in the KNS and Qamanaarsuup Sermia
regions. Whereas widespread moraine deposition in Baffin Bay occurred in
response to the 8.2 ka event (Young et al., 2020a), the GrIS retreated
inboard of the eventual historical ice limit prior to 8.2 ka at
KNS and Qamanaarsuup Sermia. Any moraine related to the 8.2 ka event that may
have existed on the landscape was overrun and destroyed by the historical
advance of the GrIS.
Retreat of the GrIS behind the modern margin during the early
Holocene
We next consider the timing and significance of when the GrIS margin crossed
the historical maximum–modern ice-margin threshold during early Holocene
deglaciation. At KNS, 10Be ages from just outboard of the historical
maximum limit indicate that following deposition of the Kapisigdlit stade
moraine at 10.24±0.36ka, the GrIS crossed the historical
maximum limit soon thereafter at 10.20±0.23ka (Fig. 2). At
Qamanaarsuup Sermia, 10Be ages from just outboard of the historical
maximum reveal that this portion of the GrIS margin crossed the historical
maximum limit at 9.29±0.18ka, ∼1kyr later
than at KNS (Fig. 3). Near Narsap Sermia (Fig. 1), 10Be ages from
outboard of the historical maximum limit indicate that the area deglaciated at
∼8.8ka (Fig. 1), and south of KNS, 10Be ages
suggest that the landscape just outboard of the historical maximum deglaciated
between ∼10.0 and 9.5 ka. Farther south, a single
10Be age suggests that the GrIS margin did not retreat behind its
modern margin until ∼7.7ka (Fig. 1). Lastly, at Saqqap
Sermia, basal radiocarbon ages suggest that deglaciation of the landscape
immediately outboard of the modern margin occurred as late as ∼5.4–4.6 ka, at least 5 kyr after deglaciation of the
landscape outboard of the modern ice margin in the KNS region (Fig. 1; Levy
et al., 2017).
At face value, there is >5kyr of spread in the timing of deglaciation
immediately outboard of the historical maximum limit in the broader KNS
region, perhaps suggestive of substantial differences in ice-margin behavior
in the early to middle Holocene (Fig. 1). When considering all of the
available ice-margin constraints, deglaciation occurred earliest in the
immediate KNS region, with deglaciation occurring later in sectors beyond
KNS. The timing of deglaciation of the landscape immediately outboard of the
historical maximum, however, is dictated by expected differences in the rate
of ice-margin retreat and differences in the magnitude of the late Holocene
readvance of the ice margin. Older 10Be ages from outboard of the
historical limit at KNS relative to adjacent ice margins could simply reflect
the earlier deglaciation of KNS compared to neighboring ice margins. In this
scenario, marine-based dynamic processes would likely drive early and rapid
deglaciation of KNS, whereas deglaciation would lag behind in adjacent
land-based ice margins where marine-based dynamical processes exert less
control on ice-margin behavior. Alternatively, the pattern of early
deglaciation at KNS with later deglaciation in adjacent margins can be
entirely explained by the magnitude of the late Holocene readvance of the ice
margin. For example, if the late Holocene readvance of KNS was of greater
magnitude than that of adjacent margins, then the KNS terminus would overrun
and rest upon a landscape that deglaciated earlier, and thus 10Be
ages from outboard of the historical moraine would be older. The greater the
magnitude of ice-margin readvance, the older the 10Be ages just
outboard of the historical moraine will be.
Additional chronological constraints that track the history of the GrIS margin
prior to and immediately after the ice margin retreated behind the
eventual historical maximum limit place the apparent asynchrony of
deglaciation across the KNS region within a broader context. At Qamanaarsuup
Sermia, 10Be ages just outboard of the historical moraine are 9.29±0.18ka; however, radiocarbon-dated lake sediments suggest that
ice remained near the historical maximum extent until ∼8.3ka (Figs. 3 and 6). Including 10Be ages from just
beyond the early Holocene moraines, all available ice-margin constraints at
Qamanaarsuup Sermia indicate that the position of the GrIS margin in this region
underwent minimal changes between ∼10.3 and ∼8.3ka, and
at least ∼2kyr of ice-margin history is represented in a
relatively restricted lateral zone on the landscape. Had the late Holocene
readvance of the Qamanaarsuup Sermia margin been slightly more extensive, the
ice margin would have overrun the early Holocene moraines currently residing
immediately outboard of the historical maximum. The resulting historical
maximum limit would abut a landscape that deglaciated just prior to
10 ka, similar to the relationship between the historical maximum
ice limit and 10Be ages outboard of this limit observed at KNS. Or,
had the late Holocene readvance been slightly less extensive, the historical
moraine would abut a landscape that likely deglaciated at ∼8.3ka, similar to the timing of deglaciation at other locations
in a broader KNS region (Fig. 1). In a similar manner, 10Be ages at
KNS indicate that the ice margin retreated behind the historical maximum limit
at 10.20±0.23ka, but ice did not continue to retreat inland and
instead remained near the current margin until ∼7.8–6.9 ka
based on concordant 10Be and in situ 14C ages from
recently exposed bedrock (Sect. 5.2). Near Narsap Sermia (Fig. 1),
10Be ages from outboard of the historical maximum limit, paired with
a basal radiocarbon age from Caribou Lake, indicate that the area deglaciated
at ∼8.8ka, but ice remained in the lake catchment and
likely near the modern margin until ∼7.5ka
(Fig. 1). Additional 10Be ages south of KNS near Deception Lake
suggest that the landscape just outboard of the historical maximum deglaciated
between ∼10.0 and 9.5 ka, but the ice margin was likely near
its present limit until ∼9.2ka based on basal radiocarbon
ages.
(a) Cosmogenic nuclide and
14C-dated proglacial–threshold lake records
constraining the behavior of the southwestern GrIS ice margin. Site
abbreviations are from Fig. 1. Using sedimentological boundaries in
proglacial–threshold lakes (organic sediments vs. silt), the orange bars
mark periods of time at each site when the GrIS was less extensive than
today (organic sediments), and blue bars marks periods of time when the ice
margin was in a position similar to today (silt). Orange dots are
traditional radiocarbon ages from the contact between basal silt and the
overlying organic sediments that constrain when the southwestern GrIS margin
retreated out of each lake's catchment in the early to middle Holocene. Blue
bullseyes are 10Be ages from immediately outboard of
the historical maximum extent in each proglacial–threshold lake drainage
catchment, and purple bullseyes are additional 10Be
ages from within ∼1km of the historical maximum extent.
Black bullseyes define our estimated timing of inland retreat of KNS based
on in situ 14C
measurements (Sect. 5.2) Note that the ice-margin constraints from
Ujarassuit Paavat (up) are traditional 14C ages from
marine bivalves reworked into the historical limit, which mark
times when the ice margin was behind the historical maximum extent (Fig. 1).
Combined, 10Be from immediately outboard of the
historical maximum limit, 10Be ages from within
∼1km of the historical maximum limit, and basal
14C ages from threshold lakes define a window when
the ice margin was in a position similar to today. Considering all
constraints, the southwestern GrIS margin likely achieved its minimum extent
after ∼5ka and was approaching its modern position as early
as ∼2ka (blue shading). (b) Mean annual temperatures at
the Greenland Ice Sheet Project 2 (GISP2) site reconstructed using gas-phase
δAr-N2 measurements (±2σ; Kobashi et al., 2017). (c) Temperature anomalies over
southwestern Greenland used in recent ISSM modeling runs (Badgeley et al.,
2020; Briner et al., 2020). (d) Corresponding modeled ice mass for the
southwestern GrIS for nine simulations using different climate
reconstructions (Briner et al., 2020); note the reversed y axis.
Additional locations with paired 10Be ages and proglacial–threshold
lake records along much of the western GrIS margin reveal that after the GrIS
margin retreated behind the position of the historical maximum extent, the ice
margin remained near this position for several hundred to thousands of years
before retreating farther inland (Fig. 17). Notably, paired 10Be
ages and radiocarbon constraints north of Jakobshavn Isbræ (Sermeq
Kujalleq) suggest that the ice margin was near its current position between
∼10 and 5.2 ka (Newspaper Lake; Cronauer et al., 2015;
Figs. 1 and 17). Closer to Jakobshavn Isbræ, paired 10Be ages
and radiocarbon constraints suggest the ice margin was in a configuration
similar to today between 7.8 and 5.5 ka (Lake Lo; Håkansson
et al., 2014; Figs. 1 and 17). In the Kangerlussuaq region, radiocarbon ages
from proglacial–threshold lakes reveal that the ice margin remained in a
configuration similar to today following initial deglaciation (Figs. 1 and 17;
Young and Briner, 2015; Lesnek et al., 2020). Unique to the Kangerlussuaq
region is proglacial lake Tasersuaq, where radiocarbon-dated sediments,
combined with maps of sub-ice topography, suggest that the GrIS margin never
retreated out of the Tasersuaq catchment, which extends only ∼1.9km behind the modern margin, during the Holocene (Lesnek et al.,
2020). Additional proglacial–threshold lake records from separate drainage
basins in the Kangerlussuaq region suggest that the ice margin retreated ∼3.7 and ∼26km inland during the Holocene (Lesnek et al., 2020)
and broadly support minimal ice-margin recession during the middle Holocene.
The combination of new 10Be ages, records from proglacial–threshold
lakes, and paired 14C–10Be measurements from KNS and
Qamanaarsuup Sermia defines a window between ∼10–7 ka when
the GrIS margin was likely near its present position. After 7 ka,
the GrIS margin retreated inland before re-approaching its current
configuration sometime in the last millennium. Considering the ice-margin
constraints from nearby Saqqap Sermia to the north (Levy et al., 2017), the
GrIS in the KNS region likely remained near its current position as late as
∼5ka. We suggest that the large range in ages constraining the
timing of deglaciation outboard of the historical moraine within the KNS
region and along the broader western GrIS margin, coupled with ice-margin
constraints from proglacial–threshold lakes and cosmogenic isotope
measurements from recently exposed bedrock, broadly defines a window between
∼10–5 ka when the GrIS margin was near its current margin
(Fig. 17). The southwestern GrIS margin reached its late Holocene maximum
extent during historical times, but records from proglacial–threshold lakes
indicate that the ice margin had advanced back to near the modern margin by
∼2ka (Fig. 17). Within a relatively narrow ∼3kyr window between ∼5–2 ka, the southwestern
GrIS margin retreated inland, achieved its minimum extent, and readvanced
towards the historical maximum extent, which likely precludes significant
inland retreat of the southwestern GrIS margin during the Holocene.
Geologic data–model comparison of ice-margin change in southwest
Greenland
Geologic reconstructions of ice-sheet change offer an ideal target for
ice-sheet modeling efforts aimed at reconstructing the geometry and ice volume
of past ice sheets through model tuning (Simpson et al., 2009; Lecavalier
et al., 2014). These reconstructions, however, also serve as an ideal test bed
for modeling efforts aimed at evaluating the sensitivity of past ice-margin
migration to climatic and oceanic influences (Briner et al., 2020). Paleo-ice-sheet modeling efforts typically rely on coarse-resolution meshes and
simplification of ice-flow approximations to achieve computational
efficiency. While this approach has enabled more simple models to assess the
sensitivity of ice-margin response to model parameters, crude climate forcings
are typically used, making it difficult to properly assess ice-margin
sensitivity to climate. Here, we further explore the deglaciation history of
the KNS region by comparing the geologic record of ice-sheet change to
recently completed model simulations of southwestern GrIS evolution through
the Holocene (Cuzzone et al., 2019; Briner et al., 2020). These Holocene ice
model simulations use the highest horizontal mesh resolution across our field
area to date and use a range of state-of-the-science gridded climate
reconstructions as the input climate (Badgeley et al., 2020; Briner et al.,
2020), thus presenting an opportunity for new insights regarding glacier
history and ice-sheet modeling. Our goals in exploring the model simulations
are to (a) assess the magnitude of recession inboard of the present margin,
(b) compare rates of retreat and timing of ice-margin change in both the model
and in the observations, and (c) explore avenues for model improvement in a
known problem area for ice-sheet modeling.
(a) Area-averaged mean annual temperature using three
different reconstructions across the ISSM model domain (Badgeley et al.,
2020; Briner et al., 2020). (b) Same as (a), but for mean annual
precipitation (Badgeley et al., 2020; Briner et al., 2020). (c) Simulated
margin position in the Sisimiut (sis)–Kangerlussuaq (kng) region through the
Holocene compared to independent observations of ice-margin position based
on dated moraines (black dots; 1σ age uncertainty; Young et al., 2020a) and the modern ice margin (“m”– orange star; modified from
Briner et al., 2020). The y axis is the distance measured from the coast. Letters
next to each simulation mark the temperature and precipitation
reconstructions from panels (a) and (b). (d) Simulated margin position in
the Nuuk (nk)–Kangiata Nunaata Sermia (kns) region through the Holocene
compared to independent observations of ice-margin position based on
deglaciation ages (see Fig. 1 for age constraints and up-fjord transect).
(e) Simulated ice mass through the Holocene across the model domain (Briner
et al., 2020); note that the y axis is reversed. (f) Simulated lateral
ice-margin extent at the time of minimum ice mass in each model run. The
year (BP) of the ice-mass minimum in each model run is listed in the legend.
Across the broader southwestern Greenland region, simulations are able to
reproduce the observed pattern of ice-margin migration through the Holocene –
early to middle Holocene retreat, inland recession, and late Holocene
readvance (Fig. 18). In particular, north of the KNS region between Sisimiut
and Kangerlussuaq, the simulated pattern and timing of ice-margin migration
generally reproduce the geologic record of ice-margin change (Briner et al.,
2020). In the KNS region, these same simulations generally depict an
ice-margin that retreats eastwards, achieves a minimum extent, and then
readvances in the late Holocene, thus sharing the broad outline of GrIS change
provided by the available geological observations (Fig. 18). However, these
simulations consistently depict ice that is too extensive compared to the
observed record of ice-margin change in the KNS region (Fig. 18). Similar to
the modeled pattern of ice-margin change, simulated ice masses in southwestern
Greenland decrease through the early Holocene and achieve a minimum value
between ∼7.6 and 6.3 ka (Fig. 18).
The better model–data fit in the Kangerlussuaq region compared to the KNS
region is likely the result of distinctly different ice-margin environments.
The Kangerlussuaq region hosts what is primarily a land-terminating sector of
the GrIS where ice-sheet behavior is dominated by surface mass balance
(Cuzzone et al., 2019; Downs et al., 2020). Thus, ice-sheet behavior in the
Kangerlussuaq region is almost entirely dictated by the climate forcings used
in the model runs (Briner et al., 2020). The KNS fjord system, however, hosted
a marine-terminating sector of the GrIS influenced by dynamical processes, but
the model setup does not include calving or submarine melting of floating
ice. Therefore, there is no dynamic mechanism with which to rapidly remove ice
from the KNS fjord system in the model. Because the ice-sheet model is built
on a high-resolution mesh, down to 2 km in the KNS region, it is able
to resolve bed features such as the existing KNS topography and fjord
bathymetry, allowing for better representation of mass transport and stress
balance than in lower-resolution models. By resolving these features, however,
ice is funneled efficiently to the ice margin, allowing the ice to persist in
the KNS fjord system in the absence of calving, which is likely responsible
for the simulated ice-margin being too extensive compared to the geologic
record. We also note that there are portions of the model domain that are
below sea level and susceptible to marine influence (Fig. S1). Throughout our simulations, relative sea level varies through
time, which could change the portions of our model domain that are marine-
vs. land-terminating. Although marine processes (e.g., submarine melting of
floating ice and calving) are not included in our simulations, we do include
grounding-line migration, and our model also simulates floating ice at outlet
glacier termini through the Holocene. It is difficult to determine how our
simulations would be impacted by including marine processes without performing
additional experiments, but areas at the ice front and immediately upstream
could be particularly affected in warmer climates coincident with the Holocene
minimum extent, as fast-flowing ice tends to maintain contact with the ocean.
Evaluating the fit between our ice-sheet model results and the
geochronological data in the KNS region must be treated cautiously as climatic
and dynamic processes may not be the only influences on model–data
mismatches. For example, the simulated GrIS retreat through the KNS region is
sensitive to how well bed topography is resolved (Cuzzone et al., 2019). In
simulations using a lower-resolution mesh (i.e., 10 km or greater),
fjord bathymetry in the KNS region is not well-resolved, and as the ice
surface lowers in response to Holocene warming, it intersects bedrock bumps
(i.e., pinning points) that would otherwise be resolved as fjords using a
higher-resolution mesh. In this scenario, ice-margin migration influenced by
bedrock sticking points due to a lower-resolution mesh might give a false
impression of good model fit to the data, but for reasons solely related to
model resolution. Regardless, simulated ice in the KNS fjord system that is
too extensive suggests that marine forcings not included in our model likely
played an important role in ice-margin retreat. While implementation and
treatment of calving in ice-sheet models continue to improve (e.g., Benn and
Åström, 2018), our results highlight the fact that inclusion of calving is
necessary towards a full understanding of ice-margin sensitivity to climate change
in the KNS region. Yet, it is difficult to properly simulate calving in
paleoclimate ice-sheet model setups that use coarse-resolution grids because
fjord systems in Greenland are typically <5km, and high-resolution
grids (1 km) are necessary to capture grounding-line migration
(Seroussi and Morlighem, 2018).
The simulated ice-margin positions in the Kangerlussuaq and KNS regions
relative to the geological constraints offer insights into model–data fit
across the entire model domain and the possible climatic conditions
influencing ice-sheet behavior through the Holocene. Among the individual
model runs of Cuzzone et al. (2019) and Briner et al. (2020), some simulated
ice masses generally achieve a minimum between ∼7.6 and
7.1 ka (simulations 1–6; Fig. 18), but there are three notable
exceptions in which the ice-mass minimum occurs later at ∼6.3ka (simulations 7–9; Fig. 18). These simulations that result
in a later ice-mass minimum also depict a minimum mass that is less extreme,
with significantly less regrowth of the ice sheet following the minimum
(Fig. 18). At the same time, these simulations with a later ice-mass minimum
depict a more subdued, although still discernible, response to the
8.2 ka cooling event relative to other simulations despite the
well-documented response of the southwestern GrIS to 8.2 ka cooling
(Figs. 17 and 18; Young et al., 2011b, 2013b, 2020a). Nonetheless, relative
sea level records from southwestern Greenland fall below modern sea level at
∼4–3 ka before later rising towards modern sea level (Long
et al., 2011; Lecavalier et al., 2014), which is interpreted to reflect the
reloading of the crust during late Holocene regrowth of the southwestern
GrIS initiating in the last few thousand years. Given all of the geologic ice-margin
constraints from across southwestern Greenland considered here, the minimum
extent of the GrIS likely occurred sometime after ∼5ka. Thus, simulations that produce an ice-mass minimum between
∼7.6 and 7.1 ka appear to be inconsistent with the
geological record. Simulations that result in an ice-mass minimum at ∼6.3ka are more compatible with the geologic record, especially
when considering the fact that these model runs simulate a subtle and broad
post-6.3 ka plateau followed by slight late Holocene regrowth
(Fig. 18).
Simulations 7 through 9 rely on a temperature history that has a muted early
Holocene warming compared to other runs, followed by peak Holocene mean annual
temperature anomalies above the 1850–2000 mean from 7 ka to
4 ka (Badgeley et al., 2020; Briner et al., 2020; Fig. 18). Run 7
relies on a precipitation history that has increased precipitation during
8 ka to 4 ka relative to the 1850–2000 CE mean, run 8
uses precipitation anomalies that are similar to the 1850–2000 CE mean, and
run 9 uses precipitation anomalies that are lower than the 1850–2000 CE
mean. Indeed, in the surface-mass-balance-dominate-dominated Kangerlussuaq region, runs 7 and 8 appear
to provide the best model–data fit of ice-margin position, while also having
an ice-mass minimum most consistent with the geologic record (Fig. 18). The
higher precipitation scenario used in run 7 is broadly supported by proxy
evidence of enhanced wintertime snowfall in southwestern Greenland and
inferred precipitation using an ice-sheet flowline model (Thomas et al., 2018;
Downs et al., 2020). Despite early Holocene warming, limited evidence
suggests that this warmth, and its effect on ice-sheet mass balance, may be
offset to some degree by increased precipitation. The timing of the minimum
inland extent of the ice margin may occur at slightly different times across
southwestern Greenland, and therefore mass does not necessarily equate to the
most retracted ice margin. However, across all model runs using different
climatologies, the simulated ice-mass minimum generally equates to a modeled
ice margin near or slightly inboard of its current position (Fig. 18).
Our results highlight the potential of proxy-informed gridded climate
reconstructions and points to where continued improvements to climate
reconstructions used in paleo-ice-sheet modeling efforts have the
biggest impact; each different climate history applied here results in a
slightly different simulated ice-sheet history. Simulated ice-sheet histories
provide a better fit to geologic constraints in the surface-mass-balance-dominated domain in
the Kangerlussuaq region, with a relatively poorer model–data fit in the KNS
region. Although some model–data mismatch occurs across our domain when
considering lateral ice-margin position, rates of GrIS mass loss inferred for
our domain (i.e., Briner et al., 2020) remain robust, as differences in
simulated ice-margin migration (i.e., slightly too fast near Kangerlussuaq and
too slow in the KNS compared to geologic constraints; Fig. 18) likely offset
each other to some degree when considered in the context of total ice-mass
loss. Ultimately, our results suggest that oceanic and dynamic processes that
are not included in our modeling effort likely play a key role in dictating
the ice-margin retreat pattern in the KNS region.
Conclusions
New 10Be ages from the KNS region, southwestern Greenland, constrain
deposition of two separate segments of the Kapisigdlit stade moraines to
10.24±0.36 and 9.57±0.38ka, indicating
that these moraines are likely not equivalent features. The older moraine is
synchronous with widespread moraine deposition in Baffin Bay at this time,
whereas the younger moraine is consistent with an additional
10Be-dated moraine in southwestern Greenland, tentatively defining a
new mode of GrIS moraine deposition at ∼9.7 to
9.6 ka. Following early Holocene moraine deposition in the KNS
forefield, the GrIS margin retreated inboard of the eventual historical
maximum extent–modern ice margin. The timing of deglaciation of the landscape
immediately outboard of the historical maximum extent in the KNS region and
along much of the southwestern GrIS margin varies by several thousand
years. Yet, additional chronological constraints provided by
proglacial–threshold lakes and cosmogenic nuclide measurements from recently
exposed bedrock surfaces constrain an interval of several thousand years during which the GrIS margin was within but near its modern position and the minimum
GrIS extent occurring sometime after ∼5ka. The variability
in 10Be ages just outboard of the historical maximum limit is likely
the result of slight variations in ice-sheet retreat and the magnitude of the
late Holocene readvance of the GrIS, rather than major differences in
ice-margin history or regional climate variability. The southwestern GrIS
margin may have advanced back to the eventual historical maximum extent as
early as ∼2ka, leaving an approximately 3kyr window for
the ice margin to achieve its minimum inland position.
Triple 10Be–14C–26Al measurements in recently
exposed bedrock fronting the modern GrIS help constrain the minimum inland
extent of the GrIS margin over multiple timescales. Our paired
26Al–10Be and 14C–10Be measurements
are unable to detect any surface burial and are consistent with constant
exposure of our sampled bedrock sites. 14C–10Be
measurements constrain the magnitude of pre-Holocene isotopic inheritance at
our bedrock sites and reveal that the period of 18th–20th century ice
cover was the only extended period that these bedrock sites became reoccupied
by the GrIS following middle Holocene
deglaciation. 26Al–10Be measurements also reveal constant
exposure of these sites and suggest that the slight amount of inherited
10Be present in a subset of our bedrock samples is due to a period
of surface exposure prior to the Last Glacial Maximum, likely during MIS 5e. In situ
14C inventories indicate that bedrock presently emerging from
beneath the GrIS in the KNS forefield was exposed during the middle
Holocene. Contemporary retreat of the GrIS has yet to expose a landscape that
remained ice-covered throughout the Holocene.
Geologic reconstructions in southwestern Greenland constrain the Holocene
behavior of the GrIS across a land-terminating region dominated by surface mass balance and an area where the ice margin retreated rapidly as the ice front
responded to dynamical changes imposed through marine influences (i.e.,
calving and/or submarine melting of floating ice). As paleo-ice-sheet models
continue to improve in terms of both the representation of processes controlling
ice-margin migration and the use of high-resolution model meshes, these geologic
reconstructions provide robust validation targets with which to benchmark and
improve models simulating the past behavior of ice sheets. Continued use of
model–data comparisons as ice-sheet modeling efforts, paleoclimate datasets,
and geologic reconstructions become more refined will improve our
understanding of past ice-sheet sensitivity to climatic and dynamic forcing
mechanisms.
Data availability
All analytical information for new cosmogenic nuclide and traditional radiocarbon measurements is listed in the tables in the Supplement. Analytical information is also available from the ICE-D: Greenland online database (http://greenland.ice-d.org/pub/26, Balco, 2020).
The supplement related to this article is available online at: https://doi.org/10.5194/cp-17-419-2021-supplement.
Author contributions
The project was conceived by NEY, JPB, and JMS. NEY, AJL, and JKC wrote the first draft of the paper with input from JPB, JAB, ABK, and JMS. All authors commented on and edited the paper. NEY, AJL, JPB, JAB, BLG, and AC completed the fieldwork. JKC led the ice-sheet modeling component, and JAB provided the regional input climatologies. RS, NEY, and JMS completed the
10Be and 26Al extraction, and NEY completed the in situ 14C extraction. MWC and SRHZ conducted final 10Be and 26Al measurements. TT and EB completed final gas-source 14C measurements. NEY, ABK, and JLL completed the final in situ 14C data reduction and interpretations. AJL and JPB isolated organic remains for traditional 14C dating.
Competing interests
The authors declare that they have no conflict of interest.
Acknowledgements
We thank the 109th Airlift Wing of the New York Air National Guard for
transport to and from Greenland, Air Greenland for helicopter support, and
CH2MHill Polar Field Services for additional logistical support. We also
thank Jean Hanley and Jeremy Frisch for help processing 10Be and 26Al samples. Accelerator mass spectrometry at CEREGE is supported by the EQUIPEX ASTER-CEREGE and ANR CARBOTRYDH (PI Edouard Bard). This is LDEO article no. 8470.
Financial support
This research has been supported by the US National Science Foundation Arctic Natural Sciences and Arctic Systems Sciences programs
(grant nos. 1417675 and 1503959 to Nicolás E. Young and Joerg M. Schaefer, grant nos. 1417783 and 1504267 to Jason P. Briner).
Review statement
This paper was edited by Alessio Rovere and reviewed by Anne Sofie Søndergaard and David Ullman.
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