9Be ratios and 10Be-fluxes (230Thxs-normalized) in central Baffin Bay sediments during the last glacial cycle: Paleoenvironmental implications

Quaternary Science Reviews 140 (2016) 142e162

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Authigenic 10Be/9Be ratios and 10Be-fluxes (230Thxs-normalized) in central Baffin Bay sediments during the last glacial cycle: Paleoenvironmental implications
s a, Laurence Nuttin b, Quentin Simon a, *, Nicolas Thouveny a, Didier L. Bourle b c, b Claude Hillaire-Marcel , Guillaume St-Onge a b c

CEREGE UM34, Aix-Marseille Universit e, CNRS, IRD, 13545 Aix en Provence, France GEOTOP Research Center, H3C3P8 Montr eal, Canada
Rimouski, Canada Institut des sciences de la mer de Rimouski (ISMER), Canada Research Chair in Marine Geology, Universit e du Qu ebec a

a r t i c l e i n f o

a b s t r a c t

Article history: Received 26 August 2015 Received in revised form 9 March 2016 Accepted 23 March 2016

Authigenic 10Be/9Be ratios and 10Be-fluxes reconstructed using the 230Thxs normalization, proxies of the cosmogenic radionuclide 10Be production rate in the atmosphere, have been measured in a sedimentary core from Baffin Bay (North Atlantic) spanning the last 136 ka BP. The normalization applied on the exchangeable (authigenic) 10Be concentrations using the authigenic 9Be isotope and 230Thxs methods yield equivalent results strongly correlated with sedimentological parameters (grain-size and mineralogy). Lower authigenic beryllium (Be) concentrations and 10Be/9Be ratios are associated with coarsegrained carbonate-rich layers, while higher authigenic Be values are related to fine-grained felsparrich sediments. This variability is due to: i) sediment composition control over beryllium-scavenging efficiency and, ii) glacial history that contributed to modify the 10Be concentration in Baffin Bay by input and boundary scavenging condition changes. Most paleo-denudation rates inferred from the 10 Be/9Be ratio vary weakly around 220 ± 76 tons.km
2.yr
1 (0.09 ± 0.03 mm.yr
1) corresponding to relatively steady weathering fluxes over the last glacial cycle except for six brief intervals characterized by sharp increases of the denudation rate. These intervals are related to ice-surging episodes coeval with Heinrich events and the last deglaciation period. An average freshwater flux of 180.6 km3.yr
1 (0.006 Sv), consistent with recent models, has been calculated in order to sustain glacially-derived 10Be inputs into Baffin Bay. It is concluded that in such environments, the authigenic 10Be measured mainly depends on climatic effects related to the glacial dynamics, which masks the 10Be production variation modulated by geomagnetic field changes. Altogether, these results challenge the simple interpretation of 10Be-concentration variation as a proxy of Interglacial/Glacial (interstadial/stadial) cycles in Arctic and sub-Arctic regions. They rather suggest the effect of higher-frequency paleoclimatic changes and local glacial dynamics on 10Be signature. © 2016 Elsevier Ltd. All rights reserved.

Keywords: 10 Be cosmogenic nuclide Authigenic 10Be/9Be ratios 10 Befluxes 230Thxs-normalization Baffin bay Late quaternary Glacial dynamics Stratigraphic markers Paleo-denudation rate Freshwater flux

1. Introduction The cosmogenic nuclide Beryllium-10 (10Be) is produced in the stratosphere (~65%) and the troposphere (~35%) by spallation reactions when highly energetic galactic cosmic ray particles interact with nitrogen and oxygen atoms (Lal and Peters, 1967; Dunai and Lifton, 2014). Its production rate is linked by a non-linear inverse relationships to the Earth and Sun magnetic fields variability

* Corresponding author. E-mail address: [email protected] (Q. Simon). http://dx.doi.org/10.1016/j.quascirev.2016.03.027 0277-3791/© 2016 Elsevier Ltd. All rights reserved.

(Elsasser et al., 1956; Lal, 1988; Beer et al., 1990; Masarik and Beer, 2009; Kovaltsov and Usoskin, 2010). After its production in the atmosphere, 10Be is quickly scavenged onto aerosol particles themselves precipitated (or thrown) within ca 1e3 years into ocean/continent reservoirs by wet or dry deposition processes (Raisbeck et al., 1981; Beer et al., 1990; Baroni et al., 2011). Previous studies have shown that -while solar activity inflects the production rate on shorter timescales- 10Be flux measured along ice and marine sediment sequences reflect large amplitude millennial scale variations of the geomagnetic dipole moment (Raisbeck et al., 1981, 1985, 1992, 2006; Yiou et al., 1985; Wagner et al., 2000; Frank et al.,

Q. Simon et al. / Quaternary Science Reviews 140 (2016) 142e162

1997; Carcaillet et al., 2003, 2004a, 2004b; Muscheler et al., 2004, nabre az et al., 2011, 2012, 2014; Valet et al., 2014). The 2005; Me exchangeable-10Be concentrations (i.e., fraction dissolved or chemically precipitated and adsorbed onto settling particles) measured in the sediments, henceforth referred to the authigenic 10 Be, are not only reflecting the atmospheric production at the time of their deposition but also depend on the scavenging rates from the overlying water column, on advected 10Be fraction due to oceanic mixing and/or on the 10Be fraction scavenged during the terrestrial/glacial/meltwater cycling of the carrier particles (Anderson et al., 1983, 1990, 1994). Hence, the normalization of authigenic 10Be concentration is a first step in order to remove these environmental components and compare 10Be records with geomagnetic variability. Two methods of normalization have been used in literature: 1) The first normalization method uses the authigenic stable 9Be isotope released by weathering of silicate rocks (von Blanckenburg et al., 2015). This method relies on the similar behavior of both isotopes once homogenized in seawater
s et al., 1989; Brown et al., 1992). The authigenic 10Be/9Be (Bourle ratio method reliably corrects for ocean/continent secondary contributions and provides robust results clearly demonstrating an inverse relationship with the geomagnetic field (HenkenMellies et al., 1990; Robinson et al., 1995; Carcaillet et al., nabre az et al., 2003, 2004a, 2004b; Thouveny et al., 2008; Me 2011, 2012, 2014). 2) The second normalization process uses the 230Th-excess method proposed by Bacon (1984) to calculate vertical fluxes of any sedimentary component deposited during the Late Quaternary (see François et al., 2004 for a review) and has been successfully used to calculate 10Be-fluxes in numerous studies (e.g., Frank et al., 1997, Frank, 2000; Christl et al., 2003, 2007, 2010). In the Arctic Ocean where most dating methods encounter serious limitations, the radioactive decay rates of the 10Be measured in marine sediments have been used to establish a Neogene chronostratigraphic framework (a priori assuming a near constant supply of 10Be) of the ACEX long sedimentary sequence (Frank et al., 2008). Besides this, the 10Be concentration variations in Arctic and sub-Arctic sediments have also been used to constrain the stratigraphy of Late Quaternary sedimentary records assuming that they are triggered by climatic cycles, the 10Be concentration during glacial periods being lower due to reduced 10Be inputs and particle fluxes within the water column (Spielhagen et al., 1997, n et al., 2009; Alexanderson et al., 2014). 2004; Selle In this study, we present new authigenic 10Be/9Be ratios and 230 Thxs-normalized 10Be-fluxes from a high-latitude marine record (HU2008-029-016PC) from central Baffin Bay spanning the last 136 ka. A primary aim of the study was to examine if the 10Be signal originating from continental inputs or advected through oceanic circulation and glacial processes, could be subtracted from total authigenic 10Be concentrations through normalization procedures in order to reveal the primary signal linked to 10Be production changes. We will thus first compare Be isotope systematics using the two available normalization methods proposed in literature, i.e., through authigenic 9Be or 230Th-excesses (230Thxs) concentrations. In the later case, we used U and Th-series isotope records from the same core (Nuttin and Hillaire-Marcel, 2015). We will then examine 10 Be-systematics vs sedimentological parameters, geomagnetic dipole moment variations, 10Be-fluxes from models and marine/ice records. Finally, we will calculate denudation rates and mean freshwater fluxes in Baffin Bay from 10Be/9Be ratios and 10Be-fluxes, and discuss paleoenvironmental implications.

143

2. Environmental setting Baffin Bay is a subpolar oceanic basin (1300 km long and 450 km wide, ~ 690,000 km2) located in the northwest North Atlantic (Fig. 1). The bay is one of the main export routes of freshwater from Greenland, the canadian Arctic and the Arctic Ocean into the North Atlantic Ocean. Its morphology consists of an elongated abyssal plain (2000e2500 m) surrounded by the continental shelves of Greenland and Baffin Island. During the glacial periods, the northeastern Laurentide Ice Sheet (LIS), the Innuitian Ice Sheet (IIS) and the Greenland Ice Sheet (GIS) formed a nearly continuous and highly dynamic ice belt surrounding Baffin Bay (Fig. 1). On the Greenland side, the GIS extended westward over the inner shelf, and as far as the shelf edge off Disko Bugt and the Uummannaq  Cofaigh et al., Trough during the Last Glacial Maximum (LGM) (O 2012, 2013; Funder et al., 2011). The LIS extended through Baffin Island, probably as far as its fjord mouths and inlets, and possibly over part of the Baffin Island shelf during glacial maxima (Margold et al., 2015). In the northern end of the bay, ice streams in the Smith Sound and Lancaster sounds were particularly large and active (England et al., 2006; Klassen and Fisher, 1988; Li et al., 2011; Stokes et al., 2012) and probably developed into an ice shelf towards the bay (Alley et al., 2010; Marcott et al., 2011). Numerous studies have demonstrated strong relationships between glacial dynamic, oceanic circulation and sedimentary processes in the bay (Aksu, 1985; Aksu and Piper, 1987; de Vernal et al., 1987; Andrews et al., 1998, 2014; Jennings et al., 2013; Simon et al., 2014; Nuttin and Hillaire-Marcel, 2015). Sedimentation occurs through two main processes and source changes: (1) glacial plumes from Greenland and Baffin Island glacial margins with large volumes of fine-grained feldspar-rich sediments transported by sediment-laden supraglacial and subglacial meltwater or nepheloid layers; and (2) TransBaffin icebergs drifting from the northern end of the bay with large volumes of coarse-grained carbonated sediments, leading to the deposition of the so-called Baffin Bay Detrital Carbonate layers (BBDC henceforth; Andrews et al., 1998; Simon et al., 2014). 3. Material and methods 3.1. Core description Core HU2008-029-016PC (70 46.14 N/-64 65.77 W; PC16 hereinafter) was raised from a 2063 m water depth, on the abyssal plain from central Baffin Bay. The 741-cm long core was retrieved using a piston corer during the 2008-029 CCGS Hudson cruise (Campbell and de Vernal, 2009). The sediment sequence is mainly composed of a succession of homogeneous dark gray to olive-gray silty clayey units and of very poorly sorted grayish-brown gravelly and sandy carbonate-rich layers (dolomite rich). These two lithofacies reflect the origin, transport and mode of deposition of the lithogenic sediments related to the ice sheet dynamics evoked above. This variability is summarized by a principal component analysis (PCA; Fig. 2) that disentangle the main compositional changes illustrated here by grain-size, density (CT number), XRF Ca/ Fe and XRD carbonate percentages along with images and CT-Scans of the study core (Fig. 3). The first principal component (PC1) accounts for 61% of the total variance and has (i) positive loadings with proxies of coarse detrital carbonate layers such as dolomite and calcite (XRD), XRF Ca, density and >63 mm %; and (ii) negative loadings with proxies related to finer sediments, feldspars (XRD) and XRF Fe and Ti (Figs. 2 and 3; Table 3). This principal component highlights the two main sedimentation processes in Baffin Bay (see Simon et al., 2014 for a discussion of sediment provenance based on mineralogical signatures). Moreover, the top of the core down to 20 cm is characterized by brown to dark brown silty muds while

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Fig. 1. General bathymetry, simplified oceanic circulation, sketch of the Paleozoic outcrops (MacLean, 1985) and paleo-Ice-Sheets (including LGM maximum ice margin extents reconstruction) of the Baffin Bay region. The location of HU2008-029-016PC sampling site is indicate by a red dots. Red arrows illustrate Atlantic “warm” waters, whereas the blue arrows represent colder Arctic waters of the Baffin Current. The simplified representations of the Greenland (green), Innuitian (blue) and Laurentide ice sheet (red) limits and major  Cofaigh et al. (2013), cLi et al. (2011), Funder et al. (2011), Dyke (2004) and England et al. (2006). (For ice stream locations (colored areas) are adapted from Margold et al. (2015), O interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

the interval between 120 and 215 cm is constituted by brownishblack to olive-black clayey muds. These two distinct lithofacies represent sediments deposited respectively during the Holocene and during the glacial phase of marine isotope stage 2 (MIS 2). The detailed sedimentological features of core PC16 result from a combination of changes in sediment delivery, transport and provenance directly related to ice margin dynamics (see Simon et al., 2012, 2014; Simon, 2013; Nuttin and Hillaire-Marcel, 2015 for additional information).

3.2. Paleomagnetic results and chronology The relative paleointensity (RPI) record previously established by a detailed paleomagnetic analysis (Simon et al., 2012) reinforced and completed preliminary paleomagnetic results obtained on another core from the same site (Thouveny, 1988). The RPI proxy (Fig. 4) is based on the normalization of the Natural Remanent Magnetization (NRM) with the Anhysteretic Remanent Magnetization (ARM) over the 25e35 mT AF demagnetization interval (NRM25e35mT/ARM25e35mT). With the exception of few problematic layers, the ARM normalizer activates the same magnetic assemblages than the NRM, and the RPI proxy presents no correlation with lithological proxies. This RPI record was correlated with RPI reference curves and stacks in order to establish the initial

chronology of the core (Simon et al., 2012). These reference records included mainly the North Atlantic relative paleointensity stack NAPIS-75 (Laj et al., 2000) and the Labrador Sea core MD95-2024 (Stoner et al., 2000). Two large inclination variations coeval with paleointensity lows allowing the recognition of two major geomagnetic excursions, i.e., the Laschamp and NorwegianGreenland-Sea excursions, and three-radiocarbon ages further supported the age model (see Simon et al., 2012 for details). This initial chronology is considered robust for the MIS 2-3-4 time interval with a high correlation between PC16 and the NAPIS-75 stack (r ¼ 0.82 between 22 and 75 ka, Fig. 4) and remains unchanged in this study. For the lower part of the core, large chronological uncertainties (low correlation coefficients with reference records) were resolved by using 13 tie points to match the PC16 RPI curve to the ODP 1063 RPI record (Channell et al., 2012) itself dated through the correlation of oxygen isotope and RPI records to calibrated reference stacks using the Match protocol (Lisiecki and Lisiecki, 2002), providing an excellent age control within MIS 5. The revised age model offers a significantly improved statistical robustness (r ¼ 0.70 between 75 and 136 ka, Fig. 4), places the PC16 core bottom (741 cm) in the MIS 6, implying a 20 ka older age at core bottom than previously estimated (Simon et al., 2012), thus refining the BBDC deposition timing (Table 1). Using this new age model, the sedimentation rate for the sedimentary sequence now

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with 20 ml MilliQ® water and conditioned with 20 ml 10.2 M HCl. The sample was then loaded onto the column and the Be fraction was collected immediately using 20 ml 10.2 M HCl for elution. The next purification step was carried out on a Dowex® 50  8 (100e200 mesh) cation-exchange resin in order to separate Boron (B) and Aluminum (Al). The resin was rinsed with 30 ml MilliQ® water and then conditioned with 30 ml 1 M HCl. After sample loading onto the column, the B and Be were eluted successively within the first 40 ml and next 120 ml of 1 M HCl eluent while the Al remained trapped within the column. After the two separation stages, Be oxy-hydroxides were precipitated at pH 8.5 from the final solution by adding NH3. The precipitate was separated by centrifugation, rinsed by re-suspension using pH 8.5 buffered MilliQ® water and centrifugated again. The purified Be oxy-hydroxides were solubilized in HNO3 and the resulting solution was transferred into a quartz crucible where it was gently evaporated to dryness at 200  C. Finally, the Be oxy-hydroxides deposit was oxidized and converted to BeO by heating at 800  C for 1 h. The BeO was then mixed with Nb powder and pressed into a cleaned Ti cathode-holder in order to condition the samples for AMS measurements. In addition to sample processing, several routine blanks and 2 replicates were measured in order to assess both cleanliness and reproducibility during the chemical extraction.

Fig. 2. Principal component analysis (PCA) of the mineralogical and grain-size dataset from PC16 (Simon et al., 2012, 2014). The loading scores for PC1 vs PC2 explain, respectively, 61 and 11.5% of the total variance. PCA analysis illustrates the two sedimentary modes in Baffin Bay.

averages 5.4 cm/ka, with some variability mostly governed by the glacial/deglacial history of surrounding lands. The mean sedimentation rates are: 1.9 cm/ka for the Holocene (0e10.6 ka), 24.9 cm/ka for the last deglacial-peak (10.7e12 ka), 5.5 cm/ka for the last glacial cycle (12e115 ka), 4.7 cm/ka for the Last Interglacial (MIS 5e) and 6.9 cm/ka for the end of MIS 6 (Fig. 4).

3.3. Sample preparation Based on the paleomagnetic record and on U and Th-series isotope records, 76 subsamples of ~1 g (dry sediment) were collected along core PC16 and processed for the Be isotope analysis at the CEREGE National Cosmogenic Nuclides Laboratory (France)
s et al. (1989) according to the chemical procedure set-up by Bourle and summarized in Carcaillet et al. (2003, 2004a, 2004b) and nabre az et al. (2011, 2012, 2014). The method is detailed here Me since the separation procedure has been modified prior to the AMS measurements. 10Be and its stable isotope 9Be were co-extracted from the authigenic phase of the sediments using 20 ml.g
1 sediment of 0.04 M hydroxylamine (NH2OHeHCl) in a 25% acetic acid leaching solution at 95 ± 5  C for 7 h. A 2 ml aliquot of the resulting leaching solution was sampled for the measurement of the natural 9 Be concentration. The remaining solution was spiked with 300 ml of a 9.8039.10
4 g.g
1 9Be-carrier before the chemical extraction in order to accurately determine 10Be sample concentrations from the measured 10Be/9Be ratios. The Be-purification was realized by chromatography in two stages. Before each separation stages, the samples were evaporated and then dissolved in ultra-pure HCl. Be oxy-hydroxides were precipitated at pH 8.5 from the solution by adding NH3. The precipitate was separated by centrifugation, dissolved in ultra-pure HCl and then processed onto a column loaded with 15 ml of specific ion-exchange resins. The iron (Fe) and manganese (Mn) were separated using a Dowex® 1  8 (100e200 mesh) anion-exchange resin. The resin was first rinsed

3.4. Measurements The natural authigenic 9Be concentrations were measured using a graphite-furnace Atomic Absorption Spectrophotometer (AAS) with a double beam correction (Thermo Scientific ICE 3400®). The standard-addition method and the addition of a constant volume of MgNO3 solution (matrix modifier) were used to eliminate the matrix effects during the absorption and to allow measurements near the detection limit. 9Be sample concentrations (Table 2) were determined from repeated absorbance measurements (4 times) performed on each of the four 100 ml aliquots of the sample solution, three of them being spiked with increasing amount of a Sharlau 9Be-carrier diluted to a known concentration (0.27e0.34  10
8 g.g
1) using HNO3 0.2%. The standard deviation of repeated absorbance measurements for each sample must be less than 3% to be accepted. After correcting for sample dilution, the authigenic 9Be sample concentrations along core PC16 vary around 1.99 ± 0.83.10
7 g.g
1. The associated uncertainties (2s) varying from 0.4 to 3.2% (average value: 1.5%) are based on the reproducibility of measurements and the least-square fitting between measured absorbance at each stages of the standard-addition method (r2 > 0.9995). The 9Be concentrations measured at the AAS are transformed in atoms as follow: 0

Authigenic½9 Beat ¼ ½9 BeAAS  msample 

NA M9 Be

(1)

where: [9Be]AAS is the concentration of natural authigenic 9Be   0 mleached
maliquot is the measured at the AAS; msample ¼ msample  mleached weight of sediment remaining after aliquot sampling; NA is the Avogadro constant (6.02214.1023 mol
1) and M9Be is the beryllium9 Molar Mass (9.0121822 g.mol
1); . The natural authigenic 10Be concentration measurements were performed at the French AMS national facility ASTER (CEREGE). 10 Be sample concentrations are calculated from the measured spiked 10Be/9Be ratios normalized to the NIST 4325 Standard Reference Material (2.79 ± 0.03  10
11; Nishiizumi et al., 2007), and are decay-corrected using the 10Be half-life of 1.387 ± 0.012 Ma (Chmeleff et al., 2010; Korschinek et al., 2010):

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Fig. 3. Beryllium isotopes results vs high-resolution physical, geochemical and mineralogical results from core PC16. Log: general simplified stratigraphy of the core; CT: CAT-Scan image of the core; HRI: high-resolution digital image; CT Number (density proxy); XRD carbonates (dolomite and calcite) cumulative percents; log(Ca/Fe): mXRF element ratio for calcium and iron measured with the ITRAX© core scanner. Grain size (%) for clay, silt fractions and coarse fractions measured at 4 cm intervals by laser diffraction. Authigenic 10Be, 9 Be and 10Be/9Be ratio are displayed in log scale. Distinct lithological facies are highlighted with colour banding. Red: uppermost brownish gray silty mud unit (Uppermost Brownish, “UB”); light green: olive-black silty to clayey mud unit (Olive Clay, “OC”); white: carbonate-rich yellowish-brown to dark-brown very poorly sorted gravelly sandy mud detrital layers (DC); dark green: olive gray to dark gray poorly sorted silty to sandy mud low carbonate detrital layers (LDC) (see text for details). Marine isotopic stages 1e6 are represented by colour boxes along the depth axis. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

Table 1 Timing and duration of the Baffin Bay Detrital Carbonate layers (BBDC). BBDC-layers 0 1 2 3 4 5 6 7 8 9 10 10b 11 a b

a

b

b

BBDC timing interval (cal ka BP)

BBDC duration (ka)

10.7e12.2 14e15 21 23.5e25 27e41 49.5e51 58 - 70? 80e90 92e102 111e117 120.5e124 126e128.5 133e136

1.5 1 <1 1.5 15 1.5 >12? 10 10 6 3.5 2.5 >3?

Numbering according to this study. Age model from this study.

decay corrected

Authigenic ½10 Beat

¼

10

Be 9 Be

 M

  ½9 Beat  mspike

 ½9 Bespike 

 NA  eðltÞ 9 M Be (2)

where: (10Be/9Be)M is the measured Be ratio; mspike and [9Be]spike are respectively the mass and the concentration of the added spike; l is the decay constant for 10Be and t is the time. The uncertainties (2s) in the measured 10Be/9Be ratios and in the calculated 10Be concentrations result from statistical and instrumental error propagation (Arnold et al., 2010) and vary from 1.3 to 3.7% (average value: 2.2%). The authigenic natural 10Be/9Be ratios are derived using equations (1) and (2):

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Table 2 AMS measurements, authigenic10Be and 9Be concentrations, authigenic10Be/9Be ratios and calculated10Be and 9Be-fluxes of core HU2008-029-016PC samples. Depth in core Age Model Sample Measured10Be/9Be (cm) (ka BP) weight (g) (10
11)* 0.5 8.5 16.5 24.5 32.5 40.5 48.5 56.5 64.5 72.5 80.5 88.5 96.5 112 128.5 136.5 152.5 168.5 184.5 200.5 208.5 218.5 232.5 248.5 256.5 264.5 271 273 275 277 279 281 283 285 287 289 291 293 295 304.5 320.5 336.5 344.5 352.5 360.5 368.5 384.5 392.5 400.5 408.5 416.5 418 420 422 437 440.5 456.5 472.5 504.5 520.5 536.5 552.5 568.5 584.5 590 600 616.5 632.5 648.5 664.5 680.5 704.5

0.80 4.85 9.68 10.83 10.98 11.41 11.91 12.13 12.76 13.39 13.93 14.31 14.68 15.43 17.61 18.30 19.36 20.43 23.04 25.41 26.15 27.05 28.54 34.55 35.79 37.57 39.56 40.13 40.70 41.33 41.89 42.34 42.79 43.35 44.02 44.69 45.37 46.04 46.71 49.00 51.65 53.84 54.87 55.89 57.14 59.01 61.05 62.07 63.09 64.88 67.06 67.61 68.16 68.70 74.35 75.62 78.58 81.21 87.96 92.17 94.76 97.35 99.94 102.62 104.15 107.33 112.56 116.84 119.84 122.96 126.77 131.75

0.994 0.963 1.000 1.000 0.990 0.983 0.988 0.961 1.074 0.932 1.014 0.992 0.971 1.086 0.970 1.003 1.023 1.012 0.985 0.987 0.965 0.978 1.015 0.964 1.012 0.950 0.956 0.955 0.986 0.955 0.983 0.967 0.986 0.996 0.970 0.998 0.990 0.999 0.981 0.977 0.985 0.974 0.975 0.980 0.986 0.998 0.975 0.976 0.969 1.002 0.992 0.986 0.973 0.978 0.956 0.979 1.002 0.974 0.954 0.973 0.973 1.000 0.993 0.990 0.994 0.974 0.991 0.990 0.989 0.990 0.974 0.976

3.011 1.976 0.797 0.144 0.218 0.060 0.038 0.089 0.408 0.526 0.159 0.101 0.061 2.647 1.033 2.048 1.251 0.682 1.280 1.934 1.219 0.252 0.025 0.032 0.075 0.048 0.028 0.023 0.037 0.333 0.862 0.968 1.560 1.941 2.126 2.227 1.951 1.689 1.465 0.993 0.190 2.280 1.675 2.183 1.607 0.141 0.116 0.148 0.086 0.110 0.119 0.071 0.098 0.103 1.438 1.888 2.212 0.283 0.066 0.788 0.116 0.058 0.067 0.075 0.051 1.051 0.401 0.496 1.361 0.127 0.079 0.290

± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ±

0.040 0.029 0.013 0.003 0.006 0.002 0.001 0.002 0.008 0.009 0.005 0.003 0.002 0.036 0.015 0.030 0.021 0.012 0.020 0.020 0.028 0.007 0.001 0.001 0.002 0.001 0.001 0.001 0.001 0.007 0.014 0.017 0.024 0.027 0.031 0.031 0.030 0.024 0.021 0.015 0.006 0.032 0.024 0.030 0.023 0.002 0.003 0.004 0.003 0.003 0.003 0.002 0.003 0.003 0.022 0.027 0.032 0.006 0.001 0.013 0.003 0.002 0.002 0.002 0.002 0.015 0.007 0.008 0.020 0.004 0.002 0.007

10 Be-fluxa (108 Authigenic decay corrected Authigenic [9Be] Authigenic 10 (1016 at/g)* [10Be] (108 at/g)* Be/9Be (10
8)* atoms.cm
2 kyr
1)

6.403 4.326 1.687 0.304 0.468 0.129 0.081 0.195 0.801 1.191 0.330 0.215 0.133 5.151 2.233 4.321 2.582 1.434 2.762 4.099 2.660 0.539 0.051 0.070 0.153 0.105 0.062 0.051 0.079 0.729 1.846 2.084 3.333 4.085 4.584 4.658 4.135 3.542 3.134 2.133 0.405 4.900 3.610 4.684 3.438 0.293 0.247 0.320 0.186 0.229 0.250 0.151 0.210 0.220 3.164 4.025 4.653 0.606 0.144 1.708 0.248 0.122 0.140 0.160 0.107 2.271 0.849 1.054 2.899 0.270 0.171 0.627

± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ±

0.085 0.064 0.027 0.007 0.012 0.004 0.003 0.005 0.016 0.021 0.010 0.006 0.004 0.070 0.033 0.064 0.044 0.026 0.043 0.060 0.038 0.015 0.002 0.002 0.005 0.003 0.002 0.002 0.003 0.016 0.031 0.037 0.051 0.058 0.069 0.067 0.064 0.051 0.046 0.033 0.014 0.070 0.052 0.067 0.052 0.005 0.008 0.010 0.006 0.007 0.007 0.005 0.006 0.006 0.051 0.061 0.070 0.014 0.003 0.031 0.007 0.004 0.004 0.005 0.004 0.035 0.015 0.019 0.045 0.009 0.005 0.015

3.269 3.060 1.930 1.229 1.048 0.898 0.740 1.028 1.090 1.508 0.973 0.966 0.810 2.507 1.554 2.150 1.940 0.949 2.385 1.777 1.855 0.709 0.692 0.590 1.323 1.178 1.372 1.233 1.372 1.426 1.366 1.293 1.531 1.860 1.880 2.081 1.878 1.790 1.682 0.990 0.936 2.003 1.834 1.977 1.693 0.812 0.851 0.920 0.861 0.925 0.811 0.621 0.832 0.847 1.517 1.809 1.844 1.331 0.793 1.521 0.839 0.798 0.774 0.841 0.705 1.302 1.068 1.123 1.476 0.771 0.782 1.338

± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ±

0.065 0.054 0.029 0.011 0.007 0.029 0.013 0.023 0.009 0.029 0.026 0.007 0.023 0.054 0.016 0.039 0.039 0.012 0.072 0.043 0.045 0.006 0.007 0.005 0.036 0.021 0.017 0.017 0.034 0.009 0.024 0.025 0.035 0.010 0.015 0.020 0.010 0.021 0.007 0.015 0.008 0.052 0.032 0.051 0.013 0.013 0.010 0.009 0.024 0.004 0.005 0.004 0.016 0.018 0.027 0.050 0.034 0.008 0.009 0.021 0.011 0.005 0.007 0.009 0.007 0.006 0.025 0.010 0.025 0.012 0.003 0.023

1.959 1.417 0.878 0.249 0.449 0.144 0.110 0.191 0.740 0.795 0.342 0.224 0.166 2.070 1.450 2.028 1.344 1.527 1.172 2.336 1.453 0.771 0.074 0.120 0.118 0.091 0.046 0.043 0.058 0.522 1.380 1.646 2.224 2.244 2.492 2.288 2.252 2.025 1.907 2.208 0.444 2.512 2.023 2.437 2.089 0.372 0.299 0.359 0.223 0.256 0.319 0.251 0.261 0.268 2.164 2.310 2.624 0.474 0.189 1.176 0.310 0.160 0.191 0.200 0.160 1.840 0.841 0.995 2.085 0.372 0.233 0.501

± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ±

0.094 0.065 0.038 0.012 0.024 0.013 0.008 0.013 0.032 0.041 0.028 0.013 0.014 0.105 0.052 0.095 0.070 0.066 0.080 0.132 0.081 0.045 0.006 0.008 0.010 0.006 0.003 0.003 0.005 0.023 0.066 0.086 0.121 0.067 0.083 0.078 0.072 0.074 0.057 0.095 0.030 0.149 0.091 0.143 0.068 0.017 0.019 0.022 0.019 0.016 0.018 0.015 0.018 0.018 0.103 0.144 0.123 0.022 0.008 0.052 0.018 0.010 0.012 0.013 0.011 0.057 0.048 0.038 0.094 0.025 0.014 0.029

9

Be-fluxa (1016 atoms.cm
2 kyr
1)

7.750 8.530 6.441 1.989 2.993 0.817 1.854 3.629

± ± ± ± ± ± ± ±

0.131 0.181 0.157 0.063 0.107 0.028 0.174 0.363

3.957 6.034 7.369 8.037 6.697 5.696 17.009 19.161

± ± ± ± ± ± ± ±

0.070 0.130 0.179 0.248 0.231 0.195 1.589 1.913

4.088 1.264 1.043 15.153 11.286 12.250 8.769 9.717 23.407 16.499 12.124

± ± ± ± ± ± ± ± ± ± ±

0.300 0.052 0.046 0.432 0.392 0.237 0.197 0.316 0.829 0.354 0.281

12.048 5.677 6.344 7.375 7.854 6.096 6.589 6.430 20.208 7.154 8.452

± ± ± ± ± ± ± ± ± ± ±

0.883 0.226 0.275 0.215 0.271 0.120 0.151 0.207 0.743 0.162 0.205

0.681 0.947 0.865 1.694

± ± ± ±

0.046 0.067 0.034 0.165

9.318 8.013 7.466 19.042

± ± ± ±

0.621 0.562 0.287 1.845

2.304 ± 0.602

7.348 21.107 1.555 24.510 18.618 17.908 7.909 2.250 1.491 2.037 1.631 1.824 5.994

± ± ± ± ± ± ± ± ± ± ± ± ±

0.165 1.496 0.046 0.973 0.837 0.648 0.145 0.092 0.066 0.084 0.098 0.097 0.929

1.849 ± 0.096 4.060 12.547 3.379 3.056 16.874 1.519 0.728 1.330 0.852

± ± ± ± ± ± ± ± ±

0.065 0.292 0.133 0.548 1.213 0.073 0.033 0.096 0.035

55.179 ± 14.407

3.945 9.795 3.595 10.022 9.457 7.557 3.895 6.234 5.143 5.859 7.540 7.362 19.424

± ± ± ± ± ± ± ± ± ± ± ± ±

0.086 0.694 0.097 0.406 0.426 0.280 0.069 0.255 0.222 0.233 0.450 0.380 3.005

7.128 ± 0.369 1.825 4.973 7.419 16.867 15.024 5.144 4.768 7.335 4.486

± ± ± ± ± ± ± ± ±

0.033 0.117 0.286 3.024 1.078 0.243 0.206 0.521 0.178

17.085 ± 0.764 1.989 ± 0.053 4.806 ± 0.174

9.797 ± 0.435 2.501 ± 0.068 5.120 ± 0.183

1.200 ± 0.041

3.433 ± 0.112

(continued on next page)

148

Q. Simon et al. / Quaternary Science Reviews 140 (2016) 142e162

Table 2 (continued ) Depth in core Age Model Sample Measured10Be/9Be (cm) (ka BP) weight (g) (10
11)* 720.5 134.25 736.5 136.17 Mean ± std. dev. Mean ± SDOM Replicate measurementsb 16.5 9.683 184.5 23.041 Blank bk1 bk2 bk3 bk4 bk5 bk6 *2-sigma uncertainties. a Normalization using b New leachates.

230

0.991 0.992

0.973 0.987

10 Be-fluxa (108 Authigenic decay corrected Authigenic [9Be] Authigenic 10 (1016 at/g)* [10Be] (108 at/g)* Be/9Be (10
8)* atoms.cm
2 kyr
1)

1.036 ± 0.022 0.977 ± 0.019 1.639 ± 1.739

1.051 ± 0.014 0.956 ± 0.017 1.331 ± 0.561

1.054 ± 0.050 1.094 ± 0.057 1.009 ± 0.859

6.697 ± 6.763

9.076 ± 7.984

0.771 ± 0.096

1.639 ± 0.202

1.331 ± 0.065

1.009 ± 0.100

6.697 ± 0.947

9.076 ± 1.118

0.790 ± 0.015 1.315 ± 0.022 (x 10
11) 0.00071 0.00120 0.00089 0.00148 0.00067 0.00053

1.708 ± 0.032 2.806 ± 0.047

1.973 ± 0.023 2.578 ± 0.062

0.870 ± 0.038 1.101 ± 0.064

6.605 ± 0.161 25.112 ± 0.844

7.631 ± 0.181 23.076 ± 0.789

Thxs (see Nuttin and Hillaire-Marcel, 2015 and Table S1 for

(3)

The measured and calculated ratios and their uncertainties are presented in Table 2. The uncertainties (2s) of the calculated authigenic 10Be/9Be ratios are derived from the propagation of both uncertainties and vary between 2.9 and 9% (average value: 5.5%). Chemistry blank ratios range from 5.3  10
15 to 1.5  10
14, which is at least 3 orders of magnitude lower than the sample 10Be/9Be ratios.

230

Be-fluxa (1016 atoms.cm
2 kyr
1)

0.488 ± 0.010 0.460 ± 0.008 0.771 ± 0.823

10  decay corrected ½10 Beat Be Authigenic 9 ¼ Be ½9 Beat

3.5. Measurements of

9

Thxs in Baffin Bay sediments

In this study, we benefit from recent U and Th-series isotope

230

Thxs results).

results from PC16 (see Table S1 in supplementary material) in order to calculate the 230Thxs-normalized 10Be (decay-corrected) and 9Be deposition rates (referred to 10Be-fluxes and 9Be-fluxes hereinafter). The Th analyses have been performed after total digestion of 1 g of sediment, while Be isotopes were extracted from 1 g of sediment after partial leaching (in order to avoid the
s et al., 1989). A constant extraction of matricial 9Be; Bourle
s et al., 1989) has been leaching efficiency of 59.7 ± 5.4% (Bourle nabre az et al. (2011) and verified based on the results from Me nabre az (2012) and can be used for calibration (see Table S2). Me Henceforth, the 10Be-fluxes calculated in this study represent minimal values. This should be considered when comparing our results against reference values, but does not prevent quantitative/ qualitative interpretations. The measurements and calculation used for determining 230Thxs are detailed in Nuttin and Hillaire-

Table 3 Correlation coefficients of Be isotopes (concentration, ratio and fluxes) and sedimentological parameters. Parameters

9

SAR 170714B (cm/kyr) XRD- Quartz (%) XRD- K-Feldspar (%) XRD- Plagioclase (%) XRD- Calcite (%) XRD- Dolomite (%) XRF- Ca/Fe XRF- K/Ti XRF- Ti/Ca XRF- Ca XRF- Fe GS- 0e2 mm (%) GS- 2e4 mm (%) GS- 4e8 mm (%) GS- 8e63 mm (%) GS- >63 mm (%) Density (g.cm-3) 9Be (at/g) 10Be (at/g) Ratio 10Be/9Be 230(Thxs)0 (dpm.g-1) 10 Be-flux (atoms.cm-2kyr-1) 9 Be-flux (atoms.cm-2 kyr-1) Denudation Rate (t.km-2 yr-1) PC1 Sedimentological parameters PC2 Sedimentological parameters

n.s. 0.51 0.55 0.54
0.62
0.51
0.73
0.66 0.71
0.71 0.69 n.s. 0.73 0.79 n.s.
0.70
0.77 1.00 0.87 0.70 0.64 0.61 n.s. n.s.
0.81 n.s.

Be (at./g)

n.s. are not significants (P < 0.001). a From Nuttin and Hillaire-Marcel (2015).

10

Be (at./g)

n.s. 0.50 0.69 0.72
0.72
0.68
0.80
0.74 0.78
0.78 0.77 n.s. 0.77 0.78 n.s.
0.73
0.76 0.87 1.00 0.94 0.67 0.74 n.s. n.s.
0.90 n.s.

10

Be/9Be ratio

n.s. 0.49 0.75 0.80
0.76
0.74
0.81
0.77 0.79
0.77 0.79 0.40 0.72 0.67 n.s.
0.63
0.73 0.70 0.94 1.00 0.57 0.81 n.s.
0.28
0.88 n.s.

230

(Thxs)0 (dpm.g
1)a

n.s. n.s. 0.39 0.42
0.48 n.s.
0.59
0.44 0.56
0.60 0.55 n.s. 0.55 0.63 n.s.
0.67
0.56 0.64 0.67 0.57 1.00 n.s.
0.43
0.41
0.60 n.s.

10

Be-flux. (atoms.cm
2.kyr
1)

n.s. 0.49 0.66 0.71
0.66
0.71
0.68
0.65 0.67
0.66 0.66 0.38 0.61 0.55 n.s.
0.52
0.67 0.61 0.74 0.81 n.s. 1.00 n.s. n.s.
0.76 n.s.

PC1 n.s.
0.68
0.77
0.84 0.80 0.89 0.95 0.85
0.94 0.92
0.91
0.58
0.86
0.82 n.s. 0.76 0.77
0.81
0.90
0.88
0.60
0.76 n.s. n.s.

Q. Simon et al. / Quaternary Science Reviews 140 (2016) 142e162

149

Fig. 4. Relative paleointensity (NRM/ARM25e35mT) tuned on RPI record of ODP Site 1063 (Channell et al., 2012) and chronostratigraphy of PC16. The RPI tie-points used in this study are represented by green dots while the red diamonds and yellow triangle are radiocarbon ages and calcite tie-points respectively (Simon et al., 2012). The incertitude in the age model is represented by shaded area. The Sedimentation accumulation rate (SAR) is derived from the age model. Baffin Bay Detrital Carbonate layers (BBDC) are represented by a gray banding and numbered according to this study. Red vertical banding represent RPI lows and large inclination swings associated to geomagnetic excursions (Simon et al., 2012). Orange vertical banding represent RPI lows associated to the Blake geomagnetic events. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

Marcel (2015). The estimated initial 230Thxs activities, recalculated in respect to the revised chronology, are extremely variable ranging from 0.118 ± 0.081 to 5.293 ± 0.212 dpm.g
1 (1.145 dpm.g
1 on average with a standard deviation of 0.977). Except for a few samples, large surplus of 230Thxs above the production from dissolved-U decay in the overlying water column points toward a sediment-focusing environment during the last glacial period. The 10Be-fluxes are calculated as follow: decay corrected

Flux½10 Be ¼

½10 Beat=g

 Z  b230

0 Axs Th
230 decay corrected

where: ½10 Beat=g

is the

10

(4)

Be concentration at the time of

deposition in atoms per gram of sediment; Z is the water depth; 230 Th production rate from the seawater 234U decays throughout the water column (2.67.10
2 dpm.m
3.y
1; François et al., 2004) and 0 Axs ¼ Axs  eð ltÞ is the scavenged Th
230 Th
230 230 Thxs initial concentration (activity) at the time of deposition. The water depth (Z) change through time was evaluated using the global sea level curve of Grant et al. (2012; Table S1). The 9Be-fluxes decay corrected are calculated by replacing the ½10 Beat=g term by [9Be]at/g. 10 The uncertainties (2s) of the calculated Be- and 9Be-fluxes are derived from the propagation of both uncertainties and vary between 1.5 and 25% (average value: 5%, Table 2). Note that solely 3 samples (273 cm, 416 cm and 504 cm) have uncertainties above 10% representing measurement outliers or major events.

b230 is the

150

Q. Simon et al. / Quaternary Science Reviews 140 (2016) 142e162

4. Results and discussion Sample concentrations, ratios and fluxes are listed in Table 2 and presented vs depth along the core photos, CT-Scan images and description in Figs. 3 and 5. Note that all results and statistical averages are reported hereafter with a ±2s uncertainty. 4.1. Authigenic

10

Be and 9Be concentrations

The authigenic 10Be (decay-corrected) concentrations vary from 0.051 ± 0.002 to 6.403 ± 0.085  108 at.g
1 (mean: 1.64,108 at.g
1; s ¼ 1.74). Such broad variability (>2 orders of magnitude) is unusual in deep marine records where 10Be concentrations usually vary by factors of 2 or 3. Comparable large amplitude variation was only found in few polar cores from the Lomonosov Ridge (central Arctic), Fram Strait and the Ross Sea (Antarctica) where 10Be concentration ranging from 0.2 to 19.5  108 at.g
1 (Eisenhauer et al., 1994; Spielhagen et al., 1997, 2004; Aldahan et al., 1997; Sjunneskog et al., 2007) were interpreted as paleoclimatic signals with high (resp. low) 10Be concentrations during Interglacial or major Interstadial (resp. Glacial) periods. Given the temporal resolution of core PC16, we can argue that the 10Be concentration varies on shorter time scales in Baffin Bay and is not directly related to Glacial/ Interglacial cycles contrary to the mechanism proposed for the Arctic Ocean. The mean 10Be concentrations in Baffin Bay are slightly lower than values from the central Arctic and Arctic/Nordic Seas of 3
4 x 108 at.g
1 and 6.7  108 at.g
1, respectively (Frank et al., 2008; Eisenhauer et al., 1994; Aldahan et al., 1997), as well as with values from lower latitude sites such as the Portuguese margin and the Golf of Papua: ~4.4  108 at.g
1 and nabre az et al., 2011, ~5.6  108 at.g
1 (Carcaillet et al., 2004b; Me 2014). A strict interpretation of the 10Be concentration in term of atmospheric fluxes between distinct sites must be avoided because of environmental parameters. However, the overall low 10Be concentration from Baffin Bay compared to these low-latitude and Arctic sites suggests relatively reduced 10Be inputs to the coring site (especially within the very low 10Be concentration interval at 232e275 cm depth). The authigenic 9Be concentration variability is about 25 times lower than that of 10Be, the 9Be concentrations varying from 0.590 ± 0.005 to 3.269 ± 0.065  1016 at.g
1 (mean: 1.33 1016 at.g
1; s ¼ 0.56). This range is still larger than the variability observed in lower-latitude marine cores, advocating for some 9Be concentration changes along core PC16. The mean 9Be concentration values in PC16 are lower than those reported for Portuguese margin and Golf of Papua sequences (2
4  1016 at.g
1), but slightly higher than values from the central Arctic (0.6e1  1016 at.g
1) and from ODP983 site (~0.25  1016 at.g
1; Knudsen et al., 2008). Despite specific interpretations related to each sites, this comparison confirms that the strong association between 9Be concentration and
s et al., 1989; Brown et al., 1992) terrigenous inputs (e.g., Bourle applies to Baffin Bay. The central and deepest part of the bay acts as a settling basin (sediment-focusing area) receiving larger 9Be inputs than deep open ocean basins namely because of its proximity to continental margins and the reduced size and shape of the bay. This is in accordance with the UeTh results from PC16 (Nuttin and Hillaire-Marcel, 2015).

2) values higher than 10
8 (Table 2, Figs. 3 and 5). 10Be concentrations and 10Be/9Be ratios of the first group are coherent with those measured in poorly sorted diamicton layers of the Ross Ice Shelf (Sjunneskog et al., 2007), in Arctic sediments of the last 350 kyr (e.g., Aldahan et al., 1997) and in Arctic rivers (10Be/9Be ratios ranging from 0.36 to 1.54  10
8; Frank et al., 2009). This supports a continental origin of the dissolved beryllium in Baffin Bay. The second group presents higher values in range with the lower limits of authigenic 10Be/9Be ratios from mid-to-low-latitude cores with similar sedimentation rates (Carcaillet et al., 2004b; nabre az et al., 2014). The average low 10Be/9Be signature corMe responds to modern seawater 10Be/9Be ratios associated with small basins experiencing high terrigenous inputs such as the Mediter
s et al., 1989; von Blanckenburg and Bouchez, ranean Sea (Bourle 2014; von Blanckenburg et al., 2015). Moreover, the 10Be/9Be ratios consistently lower than values from the North Atlantic Ocean and Labrador Sea (von Blanckenburg and Bouchez, 2014; von Blanckenburg and O'Nions, 1999) suggest too low water exchange rates in Baffin Bay to drive long-term change of the 10Be/9Be ratio in the bay. 4.3.

10

Be-fluxes (230Thxs-normalized)

The preserved vertical 10Be deposition rates (10Be-fluxes) varies from 0.68 ± 0.04 to 24.51 ± 0.97  108 at cm
2.ka
1 with a mean value of 6.7  108 at cm
2.ka
1 and a standard deviation of 6.7  108 at.cm
2.ka
1 expressing a large variability (Table 2, Figs. 5 and 6). The Holocene average 10Be-flux is estimated at 7.6 ± 0.6  108 at.cm
2.ka
1, while the 10Be-fluxes oscillated around 13.3 ± 1.9 and 1.9 ± 0.2  108 at.cm
2.ka
1 during the glacial period within the carbonate-free and carbonate-rich layers, respectively (carbonate layers are highlighted by white banding in Figs. 3 and 5, Table S4). Our computed Holocene 10Be-fluxes are similar to 10Befluxes recently modeled during the Holocene, in the Baffin Bay € and von Blanckenburg (2015) ranging from 4 to region, by Heikkila 8.8  108 at.cm
2.ka
1 (Table 4, Figs. 5 and 6). During the glacial period, 10Be-fluxes estimated for the Baffin Bay are coherent with the bulk deposition rates of 10Be from Greenland and the Arctic Ocean (ranging from 2  108 to 33  108 at.cm
2.ka
1; Eisenhauer et al., 1994; Aldahan et al., 1997; Spielhagen et al., 1997). Despite an obvious sedimentological relationship in core PC16, the 10Be-fluxes are broadly compatible with 1) modeled 10Be-production ranges (~3e27  108 at.cm
2.ka
1), 2) 10Be-fluxes from ice and marine records (~1.2e70  108 at.cm
2.ka
1) and with the globally integrated and long-term averaged 10Be-fluxes into marine sediments (9e28  108 at.cm
2.ka
1) (see Figs. 5 and 6 and Table 4 for details and references). The lowest 10Be-fluxes typical of BBDC layers are also similar to those measured in the Ice Summit record of GISP2 (2e6  108 at.cm
2.ka
1; Muscheler et al., 2004). The general agreement between 10Be-fluxes estimated here and those obtained from other sites (Eisenhauer et al., 1994; Spielhagen et al., 1997; nabre az et al., 2012) as well as with recent simulations using Me general circulation models (GCMs) confirms a rapid atmospheric € mixing of 10Be before its deposition in geological archives (Heikkila € and von Blanckenburg, 2015; Frank et al., et al., 2009, 2013; Heikkila 2009). 4.4. Comparison with relative paleointensity

4.2. Authigenic

10

Be/9Be ratios 10

9

The authigenic Be/ Be ratios ranging from 0.043 ± 0.003 to 2.624 ± 0.123  10
8 (mean: 1.01  10
8; s ¼ 0.86) disclose a variability over almost 2 orders of magnitude, very similar to those of 10Be and 9Be concentrations. The authigenic 10Be/9Be ratio population can be divided in two groups: 1) values lower than 10
8 and

In core PC16, the comparison of 10Be-proxies with the RPI record does not demonstrate any clear relationship. No systematic increases of authigenic 10Be/9Be ratios and 10Be-fluxes are observed during the large RPI lows corresponding to the Laschamp, Norwegian-Greenland Sea and post-Blake/Blake events (Fig. 6). The authigenic 10Be/9Be ratios and 10Be-fluxes do not present any

Q. Simon et al. / Quaternary Science Reviews 140 (2016) 142e162

151

Fig. 5. Authigenic 10Be/9Be ratio, 9Be- and 10Be-fluxes on depth. Red dots correspond to depths with measured 230Thxs vertical flux higher than theoretical 230Thxs vertical flux corresponding to a sediment-focusing environment. Blue dots correspond to intervals with measured 230Thxs-flux samples lower than theoretical flux and illustrate episodes of Arctic Waters outflow through Davis Strait (Nuttin and Hillaire-Marcel, 2015). The 9Be-flux is associated to weathering rate from the drainage basin. The 10Be-flux is compared to references: (1) global-average 10Be-production values from models, see Table 4 for values and references (the blue star is the mean value from Masarik and Beer, 2009); (2) Holocene € and von Blanckenburg, 2015); (3) global values measured and averaged by Monaghan meteoric 10Be-fluxes computed for the Baffin Bay area from atmospheric models (Heikkila et al., 1986; (4) 10Be-flux in GRIP/GISP2 ice core (Muscheler et al., 2005); (5) 10Be-flux in marine records representing the so-called “allowed range” from Christl et al., 2007, 2010; (6) 10 Be-flux range in deep sea sedimentary cores from the Arctic Ocean and the Norwegian Sea (Eisenhauer et al., 1994). (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

152

Q. Simon et al. / Quaternary Science Reviews 140 (2016) 142e162

significant correlation coefficients with the RPI record (Table 3). This supports the dominance of environmental imprints eas discussed below-on the 10Be deposition in Baffin Bay. The high correlation (r ¼ 0.83) between the ARM and the authigenic 10Be/9Be ratio also indicates a strong environmental imprint related to specific grain-size ranges that insure maximum Be adsorption rates (namely the very fine to fine silt range is highly correlated with the ARM: r ¼ 0.7). Such strong correlation of the 10Be-proxies with environmental parameters suggests a local signature related with the succession of climatic events.

4.5. Testing the methods

10

Be/9Be ratios and

230

Thxs-normalized

10

Be fluxes

The two existing normalization methods were independently used in former studies, but only two of them applied both methods on the same sedimentary core. The first study by Knudsen et al. (2008) obtained large uncertainties about thorium-excess (230Thxs) values precluding their use for reliable normalization. nabre az et al. (2011) demonstrated close The second study by Me agreement of the results obtained by the two methods based on a small number of samples. In PC16, the 10Be-flux variability calculated using the 230Thxs method is similar to that of the authigenic 10 Be/9Be ratio measured at the same depths (Fig. 5). The high correlation coefficient (r ¼ 0.81) and coherence between the two signals provide strong evidence that the two normalization processes yield equivalent results in a much contrasted environment where large compositional variations prevailed (Figs. 5 and 7, Table 3). A slight oceanic circulation/mixing influence over the normalization methods is however supported by higher correlation coefficients between both normalization methods when separating the rare episodes of Arctic Waters overflow through the Davis Strait (illustrated by blue dots in Figs. 5e9). These episodes have been demonstrated within core PC16 by the occurrence of short periods of 230Thxs export (230Thxs-losses, Nuttin and Hillaire-Marcel, 2015) and likely correspond to massive outflow of deep Baffin Bay water toward the Labrador Sea and the North Atlantic. During these events, the assumption of a nearly constant flux of 230Th scavenged from seawater (François et al., 2004) is possibly not fully respected, resulting in an under-correction of the 10Be concentration using the 230 Thxs normalization compared to the authigenic 9Be normalization (Fig. 7). It somehow demonstrates that normalizing authigenic 10 Be concentrations by authigenic 9Be concentrations permits an accurate correction for the total particle flux variation given their
s et al., identical origin and behavior in the water column (Bourle 1989). Both normalization processes present large variations related to lithofacies changes (Fig. 5). The normalizers (i.e., 230Thxs and 9Be) are highly correlated (Table 3) which might seem surprising given the distinct affinities (depending on particle composition) and the different scavenging residence times of both elements in the open ocean (10e50 years for Th against 500e1000 years for Be, Chase et al., 2002). The abundance of fine glaciomarine lithogenic particles presenting strong surface reactivity with both nuclides probably explained the first assertion (Roy-Barman et al., 2005; Roy-Barman, 2009). Our results also suggest quicker adsorption rates and thus shorter scavenging residence time for dissolved Be within the Baffin Bay water column probably related to the high concentration of lithogenic particles. This is consistent with previous results from the Arctic, North Atlantic basin and circum-Antartica where lower scavenging residence time for Be isotopes of about 80e200 years associated with high particles concentration and oceanic mixing have been put forth (von Blanckenburg et al., 1996; von Blanckenburg and O'Nions, 1999; von Blanckenburg and Bouchez, 2014; Frank et al., 2009).

4.6. Beryllium and sedimentological features From the authigenic Be result variations presented above, questions arise about the nature of the forcing parameters. In order to understand the links with lithofacies changes, core PC16 offers a unique opportunity because of the numerous multi-proxy results available. Table 3 presents the correlation coefficients between beryllium isotopes and the sedimentological parameters. As suggested from Figs. 3 and 6, the high correlation coefficients between PC1 and Be isotopes and ratio (>-0.8) express a strong association with sedimentological parameters. The authigenic Be values generally present significant increases at levels corresponding to fine grained feldspar-rich intervals associated with increases of clay minerals (30e50%; Simon et al., 2014), while coarse carbonate-rich sediments (30e40% dolomite and 10e15% calcite) are characterized by lower Be values (Fig. 3). This pattern is consistent with the scavenging efficiency of dissolved Be that depends on the composition and size of the particles available in the water column (Chase et al., 2002). The authigenic beryllium (10Be and 9Be) and thoriumexcess (230Thxs) distribution are strongly grain size dependent, with significant positive correlations with very fine to fine silts (2e8 mm), while clays (0e2 mm) and medium to very coarse silts (8e63 mm) does not present any significant correlations (Table 3). The high correlation of authigenic Be isotopes with the 2e8 mm fraction rather than with the clay-sized material is intriguing. Indeed an association with smaller particles is expected from their higher specific surface. We tentatively explain these results by the cohesive behavior of the clay-sized material that tends to form aggregate. These aggregates would have faster sinking rates and lower specific surface area available for the adsorption of dissolved Be explaining higher scavenging rates associated with very fine to fine silts. To the opposite, the authigenic Be concentrations are significantly anti-correlated with the coarser intervals (>63 mm, Fig. 3) related to IRD (i.e., BBDC). The strong affinity of the 10Be/9Be ratio with lithofacies changes suggests that sources of 10Be and 9Be into the bay are located on the boarding continents and that the atmospheric 10Be component directly precipitated over the bay likely represents minor amounts compared to large inputs of continental origin. Moreover, the positive slope of the linear regression between 10Be and 9Be concentrations (r ¼ 0.88, Table 3), together with higher correlations of the 10Be concentration and 10Be/9Be ratio with sedimentological parameters compared to the 9Be concentration, imply that the 10 Be/9Be ratio variations are mostly controlled by large 10Be input changes (superimposed to dilution, scavenging or extraction efficiency of both Be isotopes in the water column). This assertion is supported by relatively steady 9Be-fluxes compared to the highly variable 10Be-fluxes (Fig. 5; Table S4). This difference is likely due to distinct sources and transport processes between both isotopes. 4.7. Sources and transport processes of the

10

Be and 9Be isotopes

The 9Be isotopes are derived from the mechanical erosion and weathering (i.e., denudation) of bedrocks at the basal interface of ice streams and delivered through glacial processes into Baffin Bay. Continental 10Be (i.e., atmospheric 10Be deposited onto ice-sheets) are coming from the melting and calving of ice sheets and are released in the water column during the ice melting (from icebergs and/or freshwater discharges). A fraction of the dissolved 9Be isotopes is rapidly scavenged on the continental margins within the buoyant turbid meltwater plumes or nepheloid layers (i.e., resuspension of fine sediment particles by bottom currents) due to higher particles concentration (Bacon and Rutgers van der Loeff, 1989; Kusakabe et al., 1991), while the 10Be transported within icebergs, sea ice and freshwater plumes is exported farther from

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Fig. 6. Authigenic 10Be/9Be ratio, 10Be-flux and principal component of the PCA (PC 1) compared to relative paleointensity (RPI) vs age. PC16 10Be-flux is compared to references (see Fig. 5 and Table 4 for references). Baffin Bay Detrital Carbonate layers (BBDC) are represented by a gray banding and numbered according to Simon et al. (in prep.). Red vertical banding represent RPI lows and large inclination swings associated to geomagnetic excursions (Simon et al., 2012). Orange vertical banding represent RPI lows associated to the Blake geomagnetic events. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

the sources. The 9Be-fluxes are relatively constant with respect to lithogenic changes supporting almost constant denudation rates and/or transport processes, while 10Be-fluxes display large variations over the glacial-interglacial periods and ice rafting events. Despite this large variation, the range of 10Be-fluxes is coherent with values from ice core/marine records and recent measurements/models (Figs. 5 and 6). We impute this similarity by a smoothing of the 10Be cosmogenic nuclide signal due to the longlasting deposition/drift into the regional ice-sheets and by the thick sea ice and/or ice-shelf cover during glacial periods. The 10Befluxes calculated in Baffin Bay would thus likely represent a 10Be production signal buffered by the glacial factors controlling the Be inputs into the bay. In order to test this assumption, we calculated an average 10Be concentration from the Greenland GISP ice core and used this value to infer an amount of ice melting required to sustain the 10Be-fluxes in Baffin Bay. This led to an estimate of the total freshwater flux associated to meltwater discharge and icebergs

calving from the ice sheets that border the bay, as follows:

Total freshwater flux ¼

Flux½10 Be  ABB d  ½10 Beice

(5)

with ABB (Baffin Bay surface) ¼ 690,000 km2 and Flux½10 Be ¼ 6.7 ± 0.9 at cm
2.kyr
1 (average value for PC16). We use an average ice density value of 0.9 for simplicity and a ½10 Beice equal to 2.8 ± 1.3  104 at.g
1 corresponding to the mean 10Be concentration in GISP2 for the 3e70 ka interval (Finkel and Nishiizumi, 1997). The total freshwater fluxes obtained vary from 106.7 to 377.4 km3 yr
1 with an average value of 180.6 km3 yr
1 (0.006  106 m3 s
1 (Sv)). Although this value represents a crude estimation, it is still similar to total Greenland freshwater flux toward Baffin Bay recently computed of about 180.7 km3.yr
1 (Rignot and Kanagaratnam, 2006) or 208 ± 17 km3.yr
1 (Bamber et al., 2012). The extreme values obtained considering standard

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Table 4 10 Be-production/flux in literature (in atoms.cm
2.kyr
1). Range ( 108)

Mean values ( Notes 108)

References

10

Be-production 5.7e10 12e19 6.6 5.2e26.4 12.1 11e17 14.0 6.0 14.2 9.5 12.6 4e8.8 5.8

Global average 10Be-production (model) 60-90 N 10Be-production (model) Global average 10Be-production (model) Global average 10Be-production (analytical) Long term averaged glob. av. 10Be-prod. (analytical) Global average 10Be-production (model) Global average 10Be-production (empiric) Global average 10Be-production (analytical) Global average 10Be-production (empiric) Holocene average 10Be-flux for the central Baffin Bay area (ECHAM5-HAM general circulation model) 10 Be-flux in ice-core records 1.8e8.6 3.4 Greenland Summit (GRIP/GISP2) 3.3 Dye 3 (Greenland) 1.5e12.8 Renland Ice Core over 1931e1987 (Greenland) 1.4e10 Dronning Maud Land ice core over 1932e1988 (Antarctic) 4.0 ± 0.3 EPICA DML (Neumayer surface snow, Antarctic) 1.9 ± 0.1 EPICA DML (Kohnen deep ice core, Antarctic) 1.4 ± 0.2 Dome C (surface firn and firn core 0e12 m; Antarctica) 1.2e2.5 Epica DC between 320 and 340 kyr 0.6e3.3 1.6 Epica DC between 200 and 800 kyr 2.5 Vostok, South Pole (Antarctic) 1.4 Taylor Dome (Antarctic) 1.7 and 1.8 Concordia and Vostok over the last 60 years (Antarctic) 4.8 ± 1.6 Law Dome snow pit over the year 2001 (Antarctic) 2.0e4.5 Dome Fuji ice core over 700e1900 yr CE (Antarctica) 2.2 (2.3) GRIP snow pit (Greenland) 10Be deposition flux (1986e1990) 2.8 ECHAM5-HAM general circulation model 10Be deposition flux 10 Be-flux in marine records 8.2e31.1 16.3 ODP1063 (Bermuda Rise): 0e250 kyr BP 7.8e37.1 16.1 ODP983 (North Atlantic): 0e250 kyr BP 9e28 Globally intergrated/long-term averaged 10Be-fluxes 15e35 ODP 1063A (Bermuda Rise): 170e200 kyr BP, IB exc. 10e70 ODP 983B (North Atlantic): 170e200 kyr BP, IB exc. 5e27 ODP 925 (Ceara Rise): 0e7 Ma 5e60 Globally stacked deep-sea sediments: 0e200 kyr BP 6e7 ACEX average 10Be-flux for the past 12.3 Ma (Actic) 31.9e65.9 40.9 Portuguese Margin (20e45 kyr BP) 2e23 Four Arctic Ocean cores (F]C  S  D) 10 Be-flux in PC16 (Baffin Bay marine sediments) 0.7e24.5 6.7 ± 0.9 fluxes by facies: 6.4e8.5 7.6 ± 0.6 Facies UB (Holocene) 0.7e6 1.9 ± 0.2 Facies BBDC (coarse carbonate-rich sediments) 4e24.5 13.2 ± 1.9 Facies LDC (fine feldspar-rich sediments) 8.8e23.4 13.4 ± 1.9 Facies OC (very fine feldspar-rich sediments)

deviations associated to ½10 Beice and Flux½10 Be are also consistent with total runoff losses calculated for Greenland (Mernild et al., 2010). In order to smooth out disparities that might arise from the variation through time and uncertainty of parameters from equation (6) (i.e., source 10Be concentration, 10Be-flux and Baffin Bay surface), we integrated the average value for the entire last glacial period and obtained 2.17  107 km3 of total freshwater discharge. Again, this number is pretty close to the total iceberg flux of 1.68e2.18  107 km3 that has been modeled for Baffin Bay during the same time interval (i.e., last 120 ka; Marshal and Koutnik, 2006). These results provide convincing arguments for assessing the 10Be origin and transport mode in Baffin Bay. Therefore, despite complex relationships between granulometry and mineralogy in Baffin Bay sediments, their authigenic 10Be and 9Be signatures still reveal both some local marine control and glacial influences from surrounding lands. 4.8. Denudation rates Given the distinct origins of the

10

Be and 9Be discussed above,

Webber et al., 2007; Masarik and Beer, 1999, 2009; Kovaltsov and Usoskin, 2010 Masarik and Beer, 2009 Monaghan et al., 1986 Webber et al., 2003; Webber et al., 2007 Lal and Peters, 1967 O'Brien, 1979; O'Brien et al., 1991 Lal, 1988 € and von Blanckenburg, 2015 Heikkila

Muscheler et al., 2005; Finkel and Nishiizumi, 1997 Beer et al., 1991 Aldahan et al., 1998 Auer et al., 2009

Cauquoin et al., 2014 Cauquoin, 2013 Yiou et al., 1997 Steig et al., 1996 Baroni et al., 2011 Pedro et al., 2006 Horiuchi et al., 2008 € et al., 2008 Heikkila

Christl et al., 2007, 2010

Knudsen et al., 2008 Murayama et al., 1997 Frank et al., 1997 Frank et al., 2008 nabre az et al., 2011 Me Eisenhauer et al., 1994; Aldahan et al., 1997 This study

the 10Be/9Be ratio is sensitive to continental erosion/weathering and can be used to estimate denudation rates (von Blanckenburg et al., 2012, 2015; von Blanckenburg and Bouchez, 2014). In this view, we used the von Blanckenburg and Bouchez (2014) arithmetics in order to calculate denudation rates in the Baffin Bay area, and assess their variations during the last glacial/interglacial cycle. The 10Be/9Be ratios and 10Be-fluxes have been converted into paleodenudation rates (D) as follow:

 10

Be 9 Be



 ABB ADB

oc

10

10

Be

 Foc

Be

þ ∅del  Friv   9 ¼ 9 Be Be ∅del  D  ½9 Beparent  freac þ fdiss 9

Be

9

Be

(6)

using a global mobile fraction of 9Be ðfreac þ fdiss Þ of 0.2; a global delivery factor ∅del of 0.063 (6.3%); and assuming a mean 9Be concentration of the basin’s parent rock ½9 Beparent of 2.5 ppm, valid for large basin studies (von Blanckenburg and Bouchez, 2014; von Blanckenburg et al., 2012). The actual Baffin Bay surface (ABB) is 690,000 km2, while its drainage basin (ADB) has been estimated to

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155

Fig. 7. Authigenic 10Be/9Be ratio vs 10Be-flux. Red dots correspond to depths with measured 230Thxs vertical flux higher than theoretical 230Thxs vertical flux corresponding to a sediment-focusing environment. Blue dots correspond to intervals with measured 230Thxs-flux samples lower than theoretical flux and illustrate episodes of Arctic Waters outflow through Davis Strait (Nuttin and Hillaire-Marcel, 2015). (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

697,000 km2 by adding the western Greenland ice stream total drainage area (Rignot and Kanagaratnam, 2006) with that of the eastern Baffin Island, Devon Island and southern Ellesmere Island drainage networks flowing toward Baffin Bay (Margold et al., 2015). We used the global benthic d18O LR04 (Lisiecki and Raymo, 2005) and global sea level curve (Grant et al., 2012) in order to evaluate  changes  past ice cover and estimate the Baffin Bay surface/drainage BB basin ratio AADB variations through time (Table S1). An overall uncertainty of 20% has been retained based on the assumptions made on the equation parameters (cf. von Blanckenburg and Bouchez, 2014; Table S1). The obtained estimate of paleodenudation rates varies from 65 ± 13 to 1608 ± 321 tons.km
2.yr
1 (0.03e0.64 mm.yr
1). Throughout most of the studied core, paleo-denudation rates varied relatively little around mean value of 220 ± 76 tons.km
2.yr
1 (0.09 ± 0.03 mm.yr
1), whereas within some specific intervals, higher values are clustered around an average value of 727 ± 322 tons.km
2.yr
1 (0.29 ± 0.14 mm.yr
1; Fig. 8). Altogether, these values remain within those for denudation rates of Alpine/polar glaciers (Koppes and Montgomery, 2009; Hallet et al., 1996), but are slightly higher than values proposed for extensive ice-covered regions at similar latitude on the east coast of Greenland (Andrews et al., 1994). However, these values from Andrews et al. (1994) were presented as rough estimates and minimum ones. Thus, apart from specific intervals presenting threefold increases, the denudation rates do not vary much over the last glacial

period, despite the aggressive glacial dynamics of the Baffin Bay area. This outcome is particularly important because it means that weathering changes associated with glacial erosion and climatic cycles might have been overestimated (e.g., Willenbring and von Blanckenburg, 2010; von Blanckenburg et al., 2015). Conversely, the highest denudation rate peaks represent abrupt events mainly linked to BBDC triggering and layer deposition (Fig. 8). The amplitude of these events requires glaciological processes probably involving massive ice discharges from the northern end of the bay, as suggested by their mineralogical signatures (Simon et al., 2014). They appear in-phase with strong stadials connected to Heinrich events (i.e., H0-1, H2, H4, pre-H6, H8-9, H10; Fig. 8) and are likely the result of ice-surging episodes related to major instabilities of the Northeastern Laurentide Ice Sheet (mainly from the Lancaster Sound ice stream and from Baffin Island outlets; Simon et al., 2014). The highest denudation rate value obtained so far (1608 ± 321 ton.km
2.yr
1 at 40 ka; Fig. 8), probably related to the strongest icesurging event, is coeval with the largest Heinrich events of the last glacial period (i.e., H4; Hemming, 2004; Menviel et al., 2014). The identification of the Laschamp excursion situated just below this interval (Simon et al., 2012) further reinforces its chronological interpretation. Our new paleo-denudation rate estimates from Baffin Bay provide new evidence in favor of large-scale Northeastern Laurentide Ice Sheet instabilities, contemporaneous with major surges in the Hudson Strait area leading to the deposition of numerous Heinrich events in the North Atlantic. Moreover, the

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Fig. 8. Denudation rates derived from 10Be/9Be ratio in the Baffin Bay area during the last glacial cycle. Large and abrupt denudation rate peaks associated to ice-surging episodes within the Baffin Bay Detrital Carbonate layers (BBDC) appear in-phase with large stadial intervals corresponding to Heinrich events (Hemming, 2004). The NGRIP d18O ice core record (red curves) and synthetic Greenland record (GLT-syn, blue curve) are from www.icecores.dk and Barker et al. (2011). The July insolation calculated for 70 N is from Laskar et al. (2004). (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

most recent denudation peak suggests some intense ice streaming activity from the northern end of the bay contemporaneous with the stabilization of western Greenland ice margin retreat on the mid-shelf during the Younger Dryas cold event (Hogan et al., 2016; Sheldon et al., 2016).

5. Paleoenvironmental implications Even though interpretations resulting from the authigenic Be signature in Baffin Bay can be discussed and are likely the results of several processes (see above), we can reasonably propose that the glacial dynamics of regional ice sheets is the main internal mechanism driving 10Be inputs and delivery into the bay. In this view, Be results from PC16 bring useful information about glacial and ice margin dynamics around the bay. This is especially important since the origin, timing and limits of regional glacial variations are still very poorly constrained. The most striking features of this glacial dynamic are the variability of the GIS limits over the Greenland  Cofaigh et al., 2012, 2013; Hogan et al., 2012, continental shelf (O 2016; Dowdeswell et al., 2013; Jennings et al., 2013; Sheldon et al., 2016) and the ice streaming events from the LIS and IIS together with sea ice/ice-shelf cover variability (Li et al., 2011; Stokes et al., 2012; Simon et al., 2014: Margold et al., 2015). Glaciological processes such as ice-surging events (Fig. 8) followed by

intense ice streaming and iceberg drifting activities, likely account for peaking denudation rate values. This pattern could explain part of the deposition of long-duration BBDC layers (Fig. 10). Likewise, a tight climatic control over northern ice stream dynamics and iceberg drifting is also supported by the phasing between magnetic grain size (SIRM/kLF; Simon, 2013) and summer insolation variation (Fig. 9). The high correlation (r ¼ 0.76) between the RSL curve (Grant et al., 2012) and the cumulative inventory of measured 230 Thxs (Nuttin and Hillaire-Marcel, 2015) recalculated in respect to the revised age model clearly demonstrates that sediment fluxes toward the center of the bay were maximum during the LGM (Fig. 9). High Be concentrations, 10Be/9Be ratios and 10Be-fluxes (Figs. 5 and 9) together with sediments originating mostly from Greenland (Simon et al., 2014) also characterized this period. More generally, intervals with high authigenic Be concentrations are always associated with somewhat constant denudation rates (Fig. 8) and finer sediments originating mainly from Greenland (Simon et al., 2014). They also match extensive sea-ice cover periods that prevented iceberg drifting through the bay (Fig. 10). Such sediments originate either from (i) resuspension of fine lithogenic sediments (glacial flour) through nepheloid layer proceses, during period of intense ice-advance along ice margins, (ii) outflow of dense winter waters from continental shelves, or (iii) meltwater sediment-laden plumes during period of ice margin retreats. All

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157

Fig. 9. Paleoenvironmental interpretations of authigenic 10Be/9Be ratio variations. BBDC-layers (defined by high XRD carbonates cumulative percents) are represented by the gray vertical bars and numbered according this study. Eustatic relative sea level (RSL) is from Grant et al. (2012). SIRM/kLF is a magnetic grain size proxy (Simon, 2013). The cumulative inventory ratio compares the 230Thxs accumulation fluxes with the theoretical vertical production in the overlying water column integrated over the entire time period. Values below 1 during the Holocene indicate 230Thxs-losses while values above 1 during the glacial period suggest a long-term sediment-focusing environment (Nuttin and Hillaire-Marcel, 2015).

these scenarios imply the development of ice margins/ice streams that extend over the Greenland and Baffin Island continental shelves. Together with the inherent higher scavenging efficiency of smaller particles, the increase in particle concentration due to proximal ice margin supplies likely contributed to higher Bescavenging rates in central Baffin Bay (Fig. 10). The overall low authigenic Be concentrations, 10Be/9Be ratios and 10Be-fluxes within the BBDC layers support some Be dilution in

these IRD layers. Such a dilution is possibly associated with one or several of the following processes: (1) increases of terrigenous sediment inputs, (2) changes of sediment composition/grain-size affecting the scavenging efficiency of dissolved beryllium, (3) reduced net 10Be inputs into the center of the bay and/or (4) higher exports of 10Be by extensive sea ice/iceberg drift episodes. With the exception of the short intervals associated with major ice surging events (Fig. 8), the absence of significant increases in sedimentation

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Fig. 10. Simplified Baffin Bay paleogeography during the last glacial cycle. (a) Trans-Baffin drift characterized by IRD sediments originating from the northern ice streams, and by meltwater sediment-laden plumes from Baffin Island ice streams. (b) Extended ice margins characterized by Greenland and Baffin Island glacial flour sediments corresponding to a permanently ice-covered bay, and by extended ice margin limits over continental shelves (see section 5 for detailed interpretations of the figure). Major ice stream systems from Greenland (white), Baffin Island (blue) and Innuitian region e northern Baffin, Devon and Ellesmere islands e (yellow) are represented by arrows sizes relative to their ice drainage network. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

rates within the BBDC layers (Fig. 4), 230Thxs changes (Nuttin and Hillaire-Marcel, 2015) and/or large 9Be-flux changes (Fig. 5) favor the last three processes. The sediments that characterize the BBDC

intervals are mainly originating from the northern end of the bay supporting a more distal Greenland ice margin limit together with important ice streaming activities along the northeastern

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Laurentide and Innuitian ice sheet margins. Their signatures also support higher rates of ice drifting (Simon et al., 2014). These environmental features likely explain the lower Be concentrations within the BBDC intervals by (i) reduced Be inputs due to distal source, (ii) Be-export associated to iceberg and sea ice drifting and, (iii) lower scavenging efficiency of authigenic Be associated to coarse and carbonate particles (Fig. 10). Furthermore, despite isostatic effects over the relative sea level (RSL) around Baffin Bay, the sea level rose over the continental shelves at the time of ice margin retreats (Long et al., 2008; Simpson et al., 2009; Funder et al., 2011). During such phases of marine transgression, oceanic circulation and boundary scavenging (changes in nature and intensity, Lao et al., 1992) may have become significant processes involving large transfer of dissolved Be from the center of the bay toward the margins (Roy-Barman, 2009). In this way, the very low authigenic 10Be/9Be ratios found during the late MIS 3 (ca 40e28.5 ka BP) would account for some retreat of the GIS margin with numerous ice streaming episodes from the LIS and IIS, possibly linked to the Dansgaard-Oeschger events that followed the large glacial surge of Heinrich events 4 (Figs. 8 and 9). 6. Conclusions The authigenic 10Be cosmogenic nuclide and 9Be stable isotope data were measured along a 7.41 m sedimentary core of the subarctic basin Baffin Bay in order to attempt reconstructing the geomagnetic dipole moment variations using the 10Be/9Be ratios and 10Be-fluxes (230Thxs-normalized). The results of the two normalizations are coherent and reveal that environmental imprints directly controlled the variations of the authigenic 10Be and 9Be concentrations in Baffin Bay sediments. The contribution of the atmospheric 10Be cosmogenic nuclide production modulated by the geomagnetic field intensity is overprinted by the environmental signal. Accordingly, in such conditions, 10Be proxies do not allow to characterize geomagnetic field dipole lows linked to excursions or reversals, but rather represent a 10Be production signal buffered by the glacial factors controlling Be inputs into the bay. An average freshwater flux of 180.6 km3.yr
1 (0.006 Sv), consistent with recent model values for Baffin Bay, has been calculated in order to sustain these glacially-derived 10Be inputs. Beryllium isotopes are preferentially adsorbed on fine silicate particles associated with sediment plumes originating from lateral ice margin advances. On the opposite, they have little affinity with coarse-grained and carbonate-rich sediments (i.e., Baffin Bay Detrital Carbonate layers or BBDC) associated with iceberg and sea-ice transport originated from the North-Eastern Laurentide Ice Sheet and Innuitian Ice Sheet ice streaming events. 10Be and 9Be concentrations in Baffin Bay are thus the result of particles scavenging efficiency together with variable Be-input in response to varying transport processes, location and dynamics along the ice-margins, as well as to boundary scavenging and iceberg/sea-ice drifting. Given the distinct origins of 10Be and 9Be, from the melting of ice-sheets and the erosion/weathering of bedrocks at the basal interface of ice streams respectively, the 10Be/9Be ratio has been used to estimate denudation rates. Paleo-denudation rates calculated for Baffin Bay remains generally around mean value of 220 ± 76 tons.km
2.yr
1 (0.09 ± 0.03 mm.yr
1), thus indicating relatively steady denudation values throughout the last glacial period. Six brief intervals characterized by sharp increases of the denudation rate make an exception. These intervals are related to ice-surging episodes coeval with specific Heinrich events and the last deglaciation period, indicate northeastern Laurentide Ice Sheet instabilities contemporaneous with major surges through Hudson Strait. Our findings provide strong evidences that support the use of cosmogenic nuclide 10Be as a stratigraphic marker in the Arctic and sub-Arctic

159

regions. Yet, our results caution a straightforward use of 10Beconcentrations as a proxy of Interglacial/Glacial cycles or major Interstadial periods, and rather propose to relate the 10Be variations to higher-frequency paleoclimatic changes and glacial dynamics. Acknowledgements We thank Valery Guillou for his help during chemical sample preparation and 9Be measurements at the AAS, M. Arnold, G. Aumaître and K. Keddadouche for their valuable assistance during the measurements performed at the ASTER AMS national facility (CEREGE, Aix en Provence). Special thanks are due to Friedhelm von Blanckenburg and an anonymous reviewer for their constructive and thoughtful reviews that significantly improved this manuscript. The ASTER AMS is supported by the INSU/CNRS, the IRD and the CEAEA and by the ANR through the program “EQUIPEX Investissement d’Avenir”. This study is a contribution of the project MAGORB ANR 09 BLAN 0053 (CEREGE, IPGP, LSCE). Q. Simon thanks financial support from the EDIFICE project (ERC-2013-ADG). G. StOnge and C. Hillaire-Marcel acknowledge financial support from the Natural Science and Engineering Council of Canada (NSERC). CEREGE belongs to Observatoire des Sciences de l’Univers OSUas (Aix-Marseille Universite ). Institut Pythe Appendix A. Supplementary data Supplementary data related to this article can be found at http:// dx.doi.org/10.1016/j.quascirev.2016.03.027. References Aksu, A.E., 1985. Climatic and oceanographic changes over the past 400,000 years: evidence from deep-sea cores on Baffin Bay and Davis strait. In: Andrews, J.T. (Ed.), Quaternary Environments: Eastern Canadian Arctic, Baffin Bay and Western Greenland. Allen and Unwin, Boston, pp. 181e209. Aksu, A.E., Piper, D.J.W., 1987. Late quaternary sedimentation in Baffin bay. Can. J. Earth Sci. 24 (9), 1833e1846. €m, K., 1997. 10Be records from Aldahan, A.A., Ning, S., Possnert, G., Backman, J., Bostro sediments of the Arctic Ocean covering the past 350 ka. Mar. Geol. 144 (1e3), 147e162. Aldahan, A., Possnert, G., Johnsen, S., Clausen, H.B., Isaksson, E., Karlen, W., Hansson, M., 1998. Sixty year 10Be record from Greenland and Antarctica. Proc. Indian Acad. Sci. - Earth Planet. Sci. 107 (2), 139e147. Alexanderson, H., Backman, J., Cronin, T.M., Funder, S., Ingolfsson, O.., Jakobsson, M., € wemark, L., Mangerud, J., Ma €rz, C., Mo €ller, P., O'Regan, M., Landvik, J.Y., Lo Spielhagen, R.F., 2014. An Arctic perspective on dating Mid-Late Pleistocene environmental history. Quat. Sci. Rev. 92, 9e31. Alley, R.B., Andrews, J.T., Brigham-Grette, J., Clarke, G.K.C., Cuffey, K.M., Fitzpatrick, J.J., Funder, S., Marshall, S.J., Miller, G.H., Mitrovica, J.X., 2010. History of the Greenland ice sheet: paleoclimatic insights. Quat. Sci. Rev. 29 (15e16), 1728e1756. Anderson, R.F., Bacon, M.P., Brewer, P.G., 1983. Removal of 230Th and 231Pa at ocean margins. Earth Planet. Sci. Lett. 66, 73e90. Anderson, R.F., Lao, Y., Broecker, W.S., Trumbore, S.E., Hofmann, H.J., Wolfli, W., 1990. Boundary scavenging in the pacific ocean: a comparison of 10Be and 231Pa. Earth Planet. Sci. Lett. 96, 287e304. Anderson, R.F., Fleisher, M.Q., Biscaye, P.E., Kumar, N., Dittrich, B., Kubik, P., Suter, M., 1994. Anomalous boundary scavenging in the middle Atlantic bight: evidence from 230Th, 231Pa, 10Be and 210Pb. Deep-Sea Res. II 41, 537e561. Andrews, J.T., Milliman, J.D., Jennings, A.E., Rynes, N., Dwyer, J., 1994. Sediment thicknesses and Holocene glacial marine sedimentation rates in three east Greenland fjords (ca 68 N). J. Geol. 102, 669e683. Andrews, J.T., Kirby, M.E., Aksu, A.E., Barber, D.C., Meese, D., 1998. Late Quaternary detrital carbonate (DC-) layers in Baffin bay marine sediments (67 e74 N): correlation with Heinrich events in the north Atlantic? Quat. Sci. Rev. 17, 1125e1137. Andrews, J.T., Gibb, O.T., Jennings, A.E., Simon, Q., 2014. Variations in the provenance of sediment from ice sheets surrounding Baffin bay during MIS 2 and 3 and export to the Labrador shelf sea: site HU2008029-0008 Davis Strait. J. Quat. Sci. 29 (1), 3e13. http://dx.doi.org/10.1002/jqs.2643. ́ ́
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