Holocene oceanographic changes in SW Labrador Sea,off Newfoundland

Abstract

Benthic foraminiferal assemblages supported by selected geochemical data from three marine sediment cores collected in Placentia Bay, SE Newfoundland, are used to construct an ~13,000-year-long record of regional oceanographic changes in the SW Labrador Sea. The area is located in the boundary zone between the cold, ice-loaded Labrador Current (LC) in the north and the warm Gulf Stream (GS) waters to the south. After the Younger Dryas termination, the influence of GS-derived water increased and was further strengthened at 10.7 cal. kyr BP through enhanced northward flow of Atlantic water via the Slopewater Current. A short-term event of increased terrestrial input and water column stratification at 8.4 cal. kyr BP was likely linked to the distal drainage of glacial Lake Agassiz. After 7.3 cal. kyr BP, a stronger LC weakened the inflow of warmer subsurface waters from the GS. This may be explained by extensive meltwater release from ice sheets in Arctic Canada and is concurrent with a general shift in oceanographic conditions in the Labrador Sea region. Around 4.0 cal. kyr BP, conditions became more stable with a slight increase in salinity, indicating a decrease in meltwater transported via the LC. The Northern Hemisphere neoglacial cooling around 2.8 cal. kyr BP was characterized off SE Newfoundland by a further stabilization of the current system, dominated by the LC with some continued influx of GS water.

Keywords

benthic foraminifera Gulf Stream Labrador Current variability Lake Agassiz drainage oceanic front movement Placentia Bay

Introduction

The climate of the North Atlantic region is strongly influenced by the varying strengths of cold and warm ocean currents that feed into the subpolar gyre (SPG) (Figure 1). The western sector of the gyre, the Labrador Sea, is one of the major deep water formation sites in the North Atlantic region (Marshall and Schott, 1999), redistributing heat energy and nutrients through deep ocean circulation (Delworth and Mann, 2000; Drinkwater et al., 2013; Kuhlbrodt et al., 2007; Lohmann et al., 2008; Rhein et al., 2011; Thornalley et al., 2009). This deep convection and the involved heat loss to the atmosphere constitutes an important element of the Atlantic Meridional Overturning Circulation (AMOC), which directly influences regional climate (Thornalley et al., 2009). The AMOC plays a crucial role in redistributing heat from the tropics to the poles and significantly contributes to Northern Hemisphere climate variability (Hansen and Østerhus, 2000; Mayewski et al., 2004; Thornalley et al., 2009).

Figure 1.

(a) Map showing the major surface/subsurface currents in the North Atlantic Ocean. Blue arrows indicate cold, Polar-sourced water, while red arrows indicate warmer, Atlantic-derived water. NFL = Newfoundland, CWS = Cartwright Saddle, WGC = West Greenland Current, EGC = East Greenland Current, IC = Irminger Current. The black inset box indicates the area of B. (b) Detailed map of the surface and subsurface currents around NFL. The bathymetry is indicated by the 200-, 500- and 2500-m isobaths. Blue arrows indicate cold, Polar-sourced water, and red arrows indicate warmer, Atlantic-derived water. The locations of the three gravity cores in Placentia Bay are marked as follows: 10G = AI07-10G (blue), 12G = AI07-12G (green), 14G = AI07-14G (red).

Variability of the AMOC has been studied at different time scales (Delworth and Mann, 2000; Wu and Gordon, 2002). Interdecadal to centennial-scale variability in the AMOC is strongly forced by surface density anomalies in the North Atlantic region influencing deep water formation (Eden and Willebrand, 2001; Zhang and Vallis, 2006). Coupled atmosphere–climate models show that the strength of the Gulf Stream (GS) is closely linked to the strength of the AMOC (Joyce and Zhang, 2010). The surface and subsurface currents in the North Atlantic have seen a general cooling trend since the last Holocene Thermal Maximum (Rasmussen et al., 2011; Seidenkrantz et al., 2007, 2008; Sicre et al., 2014; Ślubowska-Woldengen et al., 2007).

The main ocean currents affecting our study area in eastern Newfoundland are the cold and relatively fresh Labrador Current (LC) and the warmer, more saline GS. Deep to intermediate water formation in the Labrador Sea (Labrador Sea Water) is directly impacted by the low salinity LC and the high salinity GS, with saline surface waters increasing deep water formation (Jones and Anderson, 2008; Katsman et al., 2004; Marshall and Schott, 1999; Schmitz and McCartney, 1993; Yashayaev, 2007). The cold LC carries much of the Arctic’s cold outflow water into the North Atlantic and the SPG (Bunker, 1976; Fratantoni and McCartney, 2010). Consequently, an increased LC flow and a weaker GS may cause increased stratification in the North Atlantic and theoretically slow down the AMOC (Jones and Anderson, 2008). Such a scenario has been hypothesized for the last interglacial period (Hillaire-Marcel et al., 2001) and the Younger Dryas (YD) (Condron and Winsor, 2012; McManus et al., 2004; Pearce et al., 2013). Interestingly, the LC seems to have increased in strength over recent decades (Sicre et al., 2014), advecting an increased amount of sea ice southwards (Weckström et al., 2013). The fluctuating strengths of these oceanic currents thus play an important role in regulating recent and past climate variations as far south as the Gulf of Maine (40°N) (Wanamaker et al., 2008).

Regional atmospheric circulation and wind strength also play important roles (Sy et al., 1997). The climate of the North Atlantic region is significantly impacted by the North Atlantic Oscillation (NAO) (Hurrell and Van Loon, 1997), which affects the flow of the LC as well as sea-surface temperatures in the North Atlantic (Flatau et al., 2003; Yashayaev, 2007). Furthermore, model simulations show that the NAO exerts some control on deep convection variability in the Labrador Sea (Guemas and Salas-Mélia, 2008).

This study presents a palaeoceanographic record of the entire YD–Holocene time period in Placentia Bay, SE Newfoundland, Canada. While previous studies have primarily focussed on sea-surface and atmospheric conditions (Jessen et al., 2011; Pearce et al., 2013, 2014a; Sicre et al., 2014; Solignac et al., 2011; Weckström et al., 2013), here we investigate the Holocene oceanographic bottom-water conditions in this area, which is situated in the boundary zone between the North Atlantic subpolar and subtropical gyres, an important site of water mass mixing (Aagaard and Carmack, 1989; Cuny et al., 2002; Van Aken et al., 2011). The bay thus provides an excellent site to capture changes in North Atlantic climate and oceanography (e.g. Pearce et al., 2013). Using high-resolution foraminiferal assemblages and elemental compositions from three marine sediment cores taken from this site, we aim to reconstruct the variability of the LC and the influence of GS on this boundary zone in the coastal waters off Newfoundland during the last 13 kyr, with primary focus on subsurface conditions of the Holocene.

Setting and modern hydrography of Placentia Bay

During the last glacial period, the seafloor of Placentia Bay was covered by the Laurentide Ice Sheet, as evidenced by drumlin fields, megascale lineations and moraines present on the floor of the bay (Brushett et al., 2007; Shaw et al., 2009, 2013). The bay was most likely free of the ice sheet by ca. 14 cal. kyr BP (Dyke, 2004), prior to the YD. The variable topography of the bay results in local differences in seabed sediments and displays varying sedimentation rates and local erosion. Surface sediments collected from different areas in the bay may thus have different ages (e.g. Pearce et al., 2013; Shaw et al., 2013; Solignac et al., 2011), making it possible to construct a longer time series when combining different cores as presented in this paper.

Placentia Bay is directly affected by the southward flowing inner branch of the cold LC (Catto et al., 1999). The properties of the two branches of the LC reflect those of their contributing sources: (1) Polar Water, entrained from the Arctic Ocean through the Canadian Arctic Archipelago and Baffin Bay via the Baffin Current; and (2) water masses originally derived from the west-flowing branch of the West Greenland Current (WGC), which includes a mixture of both polar East Greenland Current water and warmer, Atlantic-sourced Irminger Sea Water (Drinkwater, 1996) (see Figure 1a). The outer LC is composed of the WGC and the Baffin Current and carries approximately 80% of the southbound water, while the inner LC is strongly influenced by outflow from Hudson Strait that mixes with Baffin Bay water (Drinkwater, 1996; Lazier and Wright, 1993). In contrast, the GS brings warm, subtropical water northwards. It enters the southern edge of the SPG (Curry and McCartney, 2001; Flatau et al., 2003; Petrie and Anderson, 1983) and mixes, to some extent, with the LC around Newfoundland. This mixing occurs near the Grand Banks, located on the continental shelf east of Newfoundland (Figure 1). The core of the warm GS is believed to be located south of the continental slope of Newfoundland (Reverdin et al., 2003). However, via the Slopewater Current, a small bifurcation from the main GS (Keigwin and Pickart, 1999), the warmer, saline and nutrient-rich GS-derived waters also reach the southern coast of Newfoundland (Fratantoni and Pickart, 2007). Evidence for this has been shown by the presence of natural debris from Cape Hatteras (North Carolina) on beaches at the mouth of Placentia Bay (Catto et al., 2003), as well as by drifters released from St. Pierre Bank, which were observed coming into southwest Placentia Bay (Petrie and Anderson, 1983). Placentia Bay is today generally ice-free year-round (Mello and Rose, 2005); while sea ice may be present occasionally, icebergs are very rare (Catto et al., 1999).

Previous palaeoceanographic studies of Placentia Bay

Placentia Bay was found to be ideal for reconstructing past oceanic current variability by Pearce et al. (2013), who described a step-wise YD termination based on a multi-proxy analysis with special focus on fossil diatom assemblages, representing the surface-water conditions (Pearce et al., 2014a, 2014b). Other studies from Placentia Bay include Jessen et al.’s (2011) investigation of late Holocene atmospheric variability over the Labrador Sea based on the occurrence of exotic (non-native) pollen. This study found that the general atmospheric circulation pattern over the North Atlantic shifted from a dominantly southeast–northwest to a more north–south pattern around 3.0 cal. kyr BP in Newfoundland. This shift also implies a change to stronger atmospheric circulation in the winter season at that time (Jessen et al., 2011; Solignac et al., 2011).

Solignac et al. (2011), Weckström et al. (2013) and Sicre et al. (2014) investigated late Holocene sea-surface conditions around Newfoundland based on dinocyst and diatom assemblages, the IP₂₅ biomarker, and alkenone measurements, respectively. Solignac et al. (2011) concluded that the two main factors controlling sea-surface conditions off Newfoundland are large-scale variations in the LC and air–sea interactions influenced by regional-scale atmospheric circulation. Sicre et al. (2014) linked late Holocene variations in LC strength to changes in the NAO/Northern Annular Mode (NAM) through increased northwesterly winds from the Arctic during positive NAM/NAO years, and Weckström et al. (2013) found increased sea-ice conditions corresponding to the positive NAO mode.

Materials and methods

The marine sediment cores used in this study were taken during a cruise in eastern Newfoundland bays in 2007 on board the Russian R/V Akademik Ioffe. To this end, three gravity cores were collected in Placentia Bay (Figure 1, Table 1): AI07-10G, AI07-12G and AI07-14G (hereafter referred to as 10G, 12G and 14G, respectively). The cores were split in half lengthwise in the laboratory and both halves were placed in cold storage at 2°C. The ‘archive halves’ were stored intact, without any sampling, while the ‘working halves’ were sampled at regular intervals for micropalaeontological, sedimentological and geochemical analyses. The archive halves were used for x-ray fluorescence (XRF) core scanning analyses.

Table 1.

Location and measurements of the cores taken in Placentia Bay, Newfoundland.

Core name	Latitude (N)	Longitude (W)	Water depth (m)	Core length (cm)
AI07-10G	47.2389°	54.6140°	231.3	460
AI07-12G	47.1367°	54.5530°	230.1	459
AI07-14G	46.9887°	54.70078°	239	513.5

Radiocarbon dates and age models

Radiocarbon dating was carried out at the AMS ¹⁴C Dating Centre, Department of Physics and Astronomy, Aarhus University. The radiocarbon ages were calibrated using the Marine13 calibration curve (Reimer et al., 2009, 2013) and the radiocarbon calibration and age–depth modelling software Oxcal (Ramsey 2008) (Table 2, Figure 2). Sites along the Newfoundland shelf and Labrador coast indicate a reservoir age in excess of 400 years (McNeely et al., 2006), and a ΔR of 139 ± 61 years was calculated from the Marine Reservoir Correction Database created by Reimer and Reimer (2001), following the methods of Solignac et al. (2011). The age model for core 14G was presented by Pearce et al. (2014a). The chronology for core 10G is based on AMS radiocarbon measurements of nine samples of benthic foraminifera and molluscs (Table 2). An earlier version of the age model for core 12G was presented in Solignac et al. (2011) and Jessen et al. (2011) as well as in Sicre et al. (2014), where additional dates were added. The age–depth models (Figure 2) were constructed using a P_Sequence depositional model in the age–depth modelling software Oxcal 4.2 (Ramsey, 2008); the distribution of dates in the cores is also shown in Figure 3. All ages in this paper are listed as calibrated (cal) kyr BP, unless otherwise noted.

Table 2.

Radiocarbon dates from the three cores.

Core name	Lab code AAR	Sample depth (cm)	Material (species)	¹⁴C age (yr BP)	Modelled mid-depth age (cal. yr BP)	Calibrated age range (cal. year BP)	δ¹³C (‰)
14G	16810	78.5–81.5	Bivalve (Mya truncata)	9496 ± 34	10,145 ± 86	10,001–10,409	−0.49
14G	16811	155–156	Bivalve (Portlandia arctica)	9897 ± 31	10,621 ± 92	10,490–10,886	Not available
14G	16812	210.5–211.5	Bivalve (Portlandia arctica)	10,241 ± 37	10,983 ± 103	10,883–11,233	0.45
14G	16814	320–320.5	Bivalve fragment	10,687 ± 40	11,662 ± 126	11,375–12,048	−0.32
14G	16813	359.5–360	Bivalve (Portlandia arctica)	10,758 ± 39	11,905 ± 130	11,610–12,243	−0.34
14G	17647	433–436	Bivalve fragments	11,077 ± 38	12,381 ± 142	12,220–12,625	Not available
10G	15764	34–35	Mixed benthic foraminifera	1306 ± 70	901 ± 76	1055–748	−2.07 ± 0.05
10G	17060	76.5–77.5	Mixed benthic foraminifera	3993 ± 66	3940 ± 89	4129–3762	−1.7 ± 0.05
10G	15765	115.5–116.5	Mixed benthic foraminifera	4821 ± 67	5142 ± 93	5300–4955	−2.71 ± 0.05
10G	17061	146–147	Mixed benthic foraminifera	5979 ± 70	6378 ± 70	6533–6252	−2.82 ± 0.05
10G	15766	174–175	Mixed benthic foraminifera	6730 ± 69	7216 ± 67	7356–7073	−1.75 ± 0.05
10G	17062	195–196	Mixed benthic foraminifera	7199 ± 73	7592 ± 57	7705–7473	−1.99 ± 0.05
10G	15767	284–285	Gastropod (Nuculana minuta)	8072 ± 73	8527 ± 72	8675–8388	0.18 ± 0.05
10G	15768	392–393	Gastropod (Nuculana minuta)	8905 ± 70	9652 ± 90	9842–9491	0.56 ± 0.05
10G	12117	456–459	Mixed benthic foraminifera	9521 ± 86	10,522 ± 99	10,725–10,320	−2.64 ± 0.05
12G	12638	30–32	Mixed benthic foraminifera	1026 ± 36	473 ± 78	635–371	−0.96 ± 0.05
12G	16625	55–57	Mollusc	1143 ± 30	630 ± 72	683–490	−1.03 ± 0.05
12G	16626	100	Mollusc	1872 ± 35	Outlier	1461–1138	0.37
12G	16627	116–117	Mollusc	1855 ± 24	1290 ± 84	1404–1124	0.05
12G	16628	131–132	Mollusc	2163 ± 26	Outlier	1763–1411	4.7 ± 1
12G	12639	150–153	Mixed benthic foraminifera	2235 ± 70	1717 ± 113	1896–1411	−1.03 ± 0.05
12G	12640	360–363	Mixed benthic foraminifera	4487 ± 46	4511 ± 132	4770–4273	−0.66 ± 0.05
12G	12118	450–453	Mixed benthic foraminifera	5440 ± 55	5679 ± 115	5873–5490	−1.83 ± 0.05

The radiocarbon ages were calibrated using the Marine13 calibration curves (Reimer et al., 2013, 2009), with a marine reservoir age offset ΔR of 139 ± 61 years (Solignac et al., 2011).

Figure 2.

Age–depth models of cores AI07-10G (blue), 12G (green) and 14G (red). All radiocarbon dates are plotted as modelled age probability distributions (grey). The two transparent dates in 12G are outliers. The width of the coloured lines shows the 1 − σ uncertainty range of the chronology for each of the cores.

Figure 3.

The composite benthic foraminiferal record of the last ~13 kyr in Placentia Bay based on combined data from cores AI07-14G (red), AI07-10G (blue) and AI07-12G (green). The foraminiferal zonation is indicated. Foraminiferal species are shown in percentage of the entire assemblages, while benthic foraminiferal flux is shown in foraminifers per cm² per year. Coloured dots on the age axis show the calibrated radiocarbon dates shown in Table 2.

Sediment and geochemistry

The lithological description of the cores is based on visual inspection made when the cores were split; photos of the cores are given in Supplementary Figures S1–S3, available online. For core 10G, x-ray diffraction (XRD) bulk analyses were performed on selected samples (Figure 5) using a PANalytical X’pert Pro MPD with Cu–K alpha radiation and Ge monochromator (40 mA and 45 kV) at the Department of Geoscience, Aarhus University. Selected samples from 10G were analysed for grain size (Figure 5) on a Sympatec HELOS laser diffractometer with a QUIXEL module, using an R4 lens and 2-mm flow-cell at the Department of Geoscience, Aarhus University. The grain size data were evaluated using the Fraunhofer theory (Beckoff et al., 2007).

The elemental composition of the sediment (Figure 4; Supplementary Figures S4–S6, available online) was analysed using an Itrax XRF core scanner instrument at the Department of Geoscience, Aarhus University. Core 10G was measured at 0.1-cm intervals using a chromium tube; cores 12G and 14G were measured at 0.2- and 0.02-cm intervals, respectively, using a molybdenum tube. Despite differences in values between cores, patterns may be compared directly (Figure 4). Selected data have previously been shown by Solignac et al. (2011) (core 12G at lower resolution) and Pearce et al. (2013, 2014a) (core 14G). Note that the XRF analysis of core 12G has been re-done, resulting in slightly different values compared with those presented by Solignac et al. (2011), which were obtained using a Cortex instrument. For further details on core 14G, see Pearce et al. (2013, 2014a).

Figure 4.

Selected elemental records from the x-ray fluorescence (XRF) analyses of the three cores. Red = data from core 14G, green = core 12G, blue = core 10G. The elemental records are grouped and superimposed to show similarities between the cores. The data are shown in counts per second (CPS). For detailed records, see Supplementary Figures S4–S6, available online.

Foraminifera

Samples for benthic foraminiferal analyses from core AI07-10G were taken at 10-cm intervals in the entire core and at every 5 cm in the 220- to 280-cm interval to increase resolution in this interval of particular interest. The samples were soaked overnight in a solution of sodium metaphosphate (9 g per litre of water) to reduce clumping of the sediments and preserve the calcitic tests. The samples from cores AI07-12G (taken every 5 or 10 cm) and AI07-14G (taken every 5 cm) were soaked in tap water alone. The samples from all three cores were subsequently wet-sieved at 63, 100 and 1000 µm and dried in an oven set to 40°C. The 100- to 1000-µm fraction was used for the foraminiferal analysis. When possible, at least 300 individuals of benthic foraminifera were counted in each sample in 12G and 14G and at least 200 in 10G. Samples with more than 45 specimens were included in the percentage calculations, and species with a minimum of 5% relative abundance in at least one sample in one of the cores are shown in the figures. Benthic foraminiferal fluxes have been calculated based on sedimentation rates derived from the age models and the assumption of an average wet sediment density of 2.0 g cm⁻³ (Knudsen et al., 1996). The calcareous benthic foraminiferal flux was calculated as specimens per cm² per year (abbreviated as f cm⁻² yr⁻¹). Species of the genus Buccella (i.e. B. frigida, B. hannai arctica, B. calida and B. tenerrima) have been grouped because of their comparable ecological preferences and to eliminate the potential error in identification and mixing of species (Rasmussen et al., 2011; Ślubowska-Woldengen et al., 2007). The close geographical proximity of the three cores found in the same age interval has made it possible to construct one common foraminiferal assemblage (Figure 3) and XRF (Figure 4) record from all three cores. The correlation is solely based on the age models, but is supported by the similar foraminiferal assemblage and elemental composition records (see also discussion). Foraminiferal zones are defined as assemblage zones, based on the relative abundances of different benthic foraminiferal species (see Figure 3, Table 3 and section ‘Results’ for more detail). Based on the relative abundances in the composite foraminiferal assemblages, the YD and Holocene have been divided into faunal zones (A, B, C, etc.). The same zonation is used across all of the cores.

Table 3.

Age ranges and corresponding depths in each core for foraminiferal-derived faunal zones and average benthic foraminiferal flux for each zone in its corresponding core.

Zone	Range (cal. kyr BP)	12G depth (cm)	12G flux (foraminifera cm⁻² yr⁻¹)	10G depth (cm)	10G flux (foraminifera cm⁻² yr⁻¹)	14G depth (cm)	14G flux (foraminifera cm⁻² yr⁻¹)
F	0–2.8	0–232.5	46	0–61.25	53
E2	2.8–4.0	232.5–323.5	24	61.25–78.75	19
E1	4.0–7.3	323.5–459	31	78.75–179.25	22
D	7.3–10.7			179.25–460	6	30.5–167.7	77
C	10.7–11.5					167.7–294.3	48
B	11.5–12.2					294.3–406.2	12
A	12.2–12.9					406.2–513.5	3

Results

Apart from a late Holocene sandy lag deposit making up the upper ca. 30 cm of core 14G (Pearce et al., 2013, 2014a) and which is not included in the present study, core 14G covers the transition from the YD through the early Holocene (12.9–9.8 cal. kyr BP, hereafter ‘kyr BP’; average sedimentation rate: 1.6 mm yr⁻¹), while AI07-12G encompasses the late Holocene (5.6–0.2 kyr BP; average sedimentation rate: 0.8 mm yr⁻¹). Core 10G spans most of the Holocene (10.4–0 kyr BP; average sedimentation rate: 0.44 mm yr⁻¹), thus overlapping in time at the top and bottom with cores 12G and 14G, respectively. The highest temporal resolution is found in the YD and early Holocene. Together, the three cores provide a record of the entire Holocene epoch in Placentia Bay.

Lithology and geochemistry

Core photographs and XRF records of selected elements for all cores are located in Supplementary Figures S1–S6, available online; the XRF elements are also presented as a composite record in Figure 4.

AI07-14G

Below the top 30-cm sand layer, the core is characterized by olive grey to dark olive grey silt to clayey silt; mollusc remains are present in most of the core, but with highest abundances between 114 and 313.5 cm (Supplementary Figure S3, available online). Below 313.5 cm, the sediments comprise dark grey to very dark brown clay with mottled intervals and 0.5- to 1-cm-thick lenses of dark greyish brown clay. The Ti and Fe concentrations from the XRF data (Figure 4; Supplementary Figure S6, available online) show gradually decreasing values towards the top, while Ca and Sr show an opposite trend with a large maximum around 11.5 kyr BP; see Pearce et al. (2014a) for more details.

AI07-10G

The upper 3.25-m section of core 10G is composed of clayey silt; from 3.25 to 3.80 m the sediment is laminated clayey silt, while the bottom unit (3.80–4.60 m) shows some disturbance of the laminae in the clayey silt (Supplementary Figure S2, available online). Throughout the XRF record, there are generally decreasing values for Ti and Fe, albeit with a small but distinct peak at 8.4–8.25 kyr BP (Figures 4 and 5). Ca and Sr show larger variability, with a broad maximum around 9.7–9.3 kyr BP, coinciding with an increase in calcite concentrations, while there is little change in dolomite. From 7 to 2.8 kyr BP, Ca and Sr show a general decrease, followed by a slight increase during the last 3 kyr.

Figure 5.

Selected benthic foraminiferal species and geochemical parameters for the interval 10.1–7.5 cal. kyr BP in core 10G. Foraminiferal species are shown in relative percentage of the entire assemblage, benthic foraminiferal flux is shown in foraminifers per cm² per year, grain size and x-ray diffraction (XRD) are shown in relative percentage of the sample, and x-ray fluorescence (XRF) is shown in counts per second. The grey boxes highlight the low-abundance and/or low-oxygen intervals in the core.

AI07-12G

The upper 3 m of core 12G consists of silt with discontinuous 1- to 5-mm-thick laminae. The lower 1.5 m is similar in composition, but the laminations are thinner (Supplementary Figure S1, available online). The XRF for 12G shows slightly decreasing values for Fe and Ti throughout the core, with more variations in Ca and Sr, which show an increase from 2.7 to 2.2 kyr BP, and slightly decreasing values through the end of the core (Figure 4; Supplementary Figure S4, available online).

Foraminiferal assemblage zones

The foraminifera are plotted as a composite, continuous record of the foraminiferal assemblages representing the bottom-water environment through the entire period of the last ~13 kyr (Figure 3). See Table 3 for ages of zonal boundaries. The benthic foraminiferal assemblages of each of the three cores from Placentia Bay are also shown separately in Supplementary Figures S7–S9, available online. There were three samples in core 10G that fell below 45 counted foraminifera in each sample, at 370, 380 and 390 cm, and one sample in 12G at 351 cm. Only very rare (fewer than three) specimens of planktonic foraminifera were observed in a few of the samples. Details on the benthic foraminiferal flux can be seen in Figure 3, Table 3, and in Supplementary Figures S7–S9, available online, while the ecological preference of the important foraminiferal taxa is summarized in Table 4.

Table 4.

Environmental preference of foraminiferal species mentioned in text and figures.

Species	Arctic water	Atlantic water	Primary productivity	Low salinity	Full-saline conditions	Bottom current activity	Notes	References
Bolivinellina pseudopunctata Höglund, 1947		X	X				Found in areas of Atlantic water off West Greenland; often blooms in connection to high food availability	Knudsen et al. (2008), Lloyd et al. (2005)
Buccella frigida (Cushman, 1952)	X		X				Found around Greenland and northern Iceland in areas influenced by Arctic water; may tolerate reduced salinities	Hald and Steinsund (1996), Jennings et al. (2011), Knudsen et al. (2008), Korsun and Hald (1998), Steinsund (1994)
Buccella frigida var. calida (Cushman and Cole, 1930)			X				See Buccella frigida
Buccella tenerrima (Bandy, 1950)	X		X				See Buccella frigida
Cassidulina reniforme (Nørvang, 1945)	X				X		Arctic species found in connection to stable bottom-water conditions	Knudsen et al. (1996), Korsun and Hald (2000), Lloyd et al. (2005), Scott et al. (2008a)
Cibicides lobatulus (Walker and Jacob, 1798)					X	X	Indicates high bottom current activity and found in coarse, sandy sediments	Murray (1991), Polyak et al. (2002), Rytter et al. (2002), Steinsund (1994)
Elphidium albiumbilicatum (Weiss, 1954)				X			Tolerates low salinities	Alve (1995), Austin and Sejrup (1994)
Elphidium bartletti Cushman, 1933	X			X			Arctic species; found in near-shore areas influenced by freshwater runoff in the Arctic	Seidenkrantz and Knudsen (1994), Polyak et al. (2002)
Elphidium excavatum forma clavata Cushman, 1941	X						Opportunistic species; known to tolerate unstable conditions (reduced salinity, high turbidity, cold water)	Knudsen et al. (1996), Korsun and Hald (2000)
Epistominella arctica Green, 1959			X				Associated with high food availability, phytoplankton blooms and deep Arctic water	Scott et al. (2008a) (as Stetsonia arctica), Wollenburg and Kuhnt (2000)
Glabratella arctica Scott and Vilks, 1991			X				Associated with high food availability and sea ice in the Arctic	Scott and Vilks (1991)
Globobulimina auriculata arctica Höglund, 1947			X				Thrives in high-productivity environments; very tolerant to low-oxygen conditions in bottom waters	Alve and Bernhard (1995), Corliss (1991)
Islandiella helenae Feyling-Hansen and Buzas, 1976	X						Found around Greenland, Iceland and the Arctic Ocean; thrives in areas with seasonal sea-ice cover	Jennings et al. (2004), Jennings and Helgadottir (1994), Korsun and Hald (1998), Seidenkrantz (2013), Steinsund (1994)
Islandiella norcrossi (Cushman, 1933)		X			X		Generally considered Arctic to subarctic, but found associated with Atlantic water in Disko Bugt, and in Svalbard fjords with transformed Atlantic water and high-saline bottom water	Hald and Korsun (1997), Hald and Steinsund (1996), Korsun and Hald (1998), Perner et al. (2011), Polyak et al. (2002), Steinsund (1994)
Nonionellina labradorica (Dawson, 1860)			X		X		Common in oceanic frontal zones; preference for fresh phytodetritus as food source	Feyling-Hanssen et al. (1971), Jennings et al. (2004), Rytter et al. (2002)
Quinqueloculina seminulum Linné, 1758		X			X		Associated with saline, warmer water; may be linked to Atlantic water	Wilkinson (2005)

Cassidulina reniforme and Elphidium excavatum forma clavata dominate throughout the records; the zones are distinguished by the occurrence of other associated species, denoting changes in palaeoceanographic conditions (next section). Foraminifera from Zones A, B, C and part of D in core 14G have been previously published by Pearce et al. (2013); here, we present faunal zones based on the combined foraminiferal data from all three cores. There was indication of some etching and dissolution of calcareous tests in the assemblages, possibly linked to high levels of CO₂ and/or dense bottom water (Aksu, 1983; De Vernal et al., 1992; Steinsund and Hald, 1994), or low levels of oxygen at the seafloor (Aksu, 1983). In general, there was more evidence of etching during intervals of lower benthic foraminiferal flux:

Zone A (12.9–12.2 kyr BP; core 14G) is defined by high relative abundances of C. reniforme, which decreases near the end of the zone, and E. excavatum f. clavata, which increases through the zone. The averaged benthic foraminiferal flux for this zone was very low, at only 3 foraminifera per cm² per year.

Zone B (12.2–11.5 kyr BP; core 14G) has high abundances of E. excavatum f. clavata and C. reniforme, albeit with lower abundances of C. reniforme than in Zone A. Discrete peaks in Buccella spp. and Glabratella arctica are seen. The averaged benthic foraminiferal flux continued to be low, at 12 foraminifera per cm² per year.

Zone C (11.5–10.7 kyr BP; core 14G) is dominated by E. excavatum f. clavata and C. reniforme, but particularly characterized by the increased abundance of E. albiumbilicatum and N. labradorica. The averaged benthic foraminiferal flux was 48 foraminifera per cm² per year.

Zone D (10.7–7.3 kyr BP; cores 14G and 10G) is dominated by C. reniforme and a sharp decrease in E. excavatum f. clavata and defined by an increase in N. labradorica and the highest abundance in the record of C. lobatulus and Q. seminulum. I. helenae gradually increases in the zone. Bolivinellina pseudopunctata peaks in a single sample around 9.2 kyr BP, while Globobulimina auriculata arctica peaks in a short interval from 8.4 to 8.25 kyr BP. There is a notable interval of very low foraminiferal abundances prior to the peak in B. pseudopunctata (9.7–9.3 kyr BP). The averaged benthic foraminiferal fluxes for cores 14G and 10G, respectively, were 77 and 6 foraminifera per cm² per year.

Zone E (7.3–2.8 kyr BP; cores 10G and 12G) is a dynamic zone characterized by high relative abundance of E. bartletti that are being substituted by the gradually increasing I. helenae, Buccella spp. and I. norcrossi. The zone is still dominated by C. reniforme and E. excavatum f. clavata. The zone is subdivided into two subzones: E1 (7.3–4.0 kyr BP) and E2 (4.0–2.8 kyr BP), and the boundary is defined at the drop in E. bartletti and N. labradorica and the peak in the Buccella group. The averaged benthic foraminiferal fluxes for cores 10G and 12G, respectively, were 22 and 31 foraminifera per cm² per year for Subzone E1, and 19 and 24 foraminifera per cm² per year for Subzone E2.

Zone F (2.8–0 kyr BP; cores 10G and 12G) is dominated by C. reniforme and E. excavatum f. clavata, with steady abundances of I. helenae, Buccella spp., E. bartletti, N. labradorica and C. lobatulus. The averaged benthic foraminiferal fluxes for cores 10G and 12G, respectively, were 53 and 46 foraminifera per cm² per year.

Palaeoenvironmental interpretation of foraminiferal zones

The foraminifera in the YD interval, supporting the two-part division (Zones A and B) linked to northward movement of the GS–LC frontal zone in the mid YD, have previously been described in detail by Pearce et al. (2014a).

In Zone C (11.5–10.7 kyr BP), corresponding to the early Holocene, subsurface waters were primarily characterized by the presence of LC water, as indicated by the high abundances of C. reniforme and E. excavatum f. clavata, as was the case for the YD period (Pearce et al., 2014a). However, the presence of N. labradorica, which is a species that has often been described in relation to oceanic frontal zones (Steinsund, 1994) or linked to relatively warm water in the western Labrador Sea and off Nova Scotia (Scott et al., 2008b), suggests a northward movement of the GS–LC frontal zone, compared with conditions during the YD (Pearce et al., 2014a). An increased influence of the warmer, nutrient-rich GS-derived waters may also be responsible for the high foraminiferal productivity indicated by the foraminiferal fluxes in core 14G (Figure 3). A relatively productive environment is supported by the comparatively high Ca and Sr concentrations (Zielinski et al., 2015) in Zone C (Supplementary Figure S6, available online).

Zone D (10.7–7.3 kyr BP) demonstrates the highest bottom-water energy level in the record (indicated by C. lobatulus), providing evidence for changes in bottom-water turbulence. This is supported by the upwards coarsening grain size (Figure 5), indicating increasing bottom current activity through the zone. This may be linked to an increasingly open, ice-free environment and growing influence of the GS, also indicated by the presence of warmer species such as Q. seminulum and N. labradorica. An interval with very low numbers of foraminifera (<45 per sample), from 9.7 to 9.3 kyr BP, indicating hostile bottom-water conditions and possibly carbonate dissolution, followed shortly after by a distinct peak of Bolivinellina pseudopunctata in a single sample around 9.2 kyr BP, indicating a short-lived high-productivity event. This species is commonly associated with high primary productivity and may also tolerate reduced oxygen concentrations of the bottom waters (Gustafsson and Nordberg, 2001; Patterson et al., 2000).

A pronounced peak in the low-oxygen tolerating species Globobulimina auriculata arctica (Alve and Bernhard, 1995; Corliss, 1991) is found in a short interval from 8.4 to 8.25 kyr BP. The peak in G. auriculata arctica coincides with a relatively small but distinct maximum in both Ti and Fe (Figures 4 and 5), indicating an increase in terrestrial input into Placentia Bay. Such increased terrestrial influx may also have led to the minor increase in the silt fraction. The peak in G. auriculata arctica indicates a reduction in the ventilation of bottom waters at that time, possibly caused by increased stratification of the water column. Nevertheless, low-oxygen conditions may only have been seasonal or may only have occurred some years, as the general foraminiferal assemblages, especially the presence of C. lobatulus, indicate that some bottom current activity continued throughout the interval.

At the start of Zone E (7.3–2.8 kyr BP), the high abundances of E. bartletti combined with the continued significance of C. reniforme and near-disappearance of C. lobatulus suggest a southward shift in the oceanic front between the LC and the GS and thus increased importance of LC water. The high abundance of E. bartletti in Subzone E1 (7.3–4.0 kyr BP) suggests slightly reduced bottom-water salinities, possibly due to more meltwater in the LC, as this species has been described as abundant off Arctic river outflow (Polyak et al., 2002). This is supported by the relatively high Ti and Fe values, indicating a higher influx of land-derived material (Wei et al., 2003) (Figure 4; Supplementary Figure S5, available online). Both E. bartletti and terrestrially derived elements decrease over time, suggesting gradually increasing salinities, as also supported by the high relative abundance of C. reniforme and the appearance of I. norcrossi after ca. 6 kyr BP. In Subzone E2 (4.0–2.8 kyr BP), E. bartletti decreased significantly and fluctuating abundances of Buccella spp., I. helenae and generally variable foraminiferal flux indicate overall less Arctic meltwater influence, but somewhat varying conditions with peaks of higher productivity, possibly linked to local changes in nutrient conditions or even small fluctuations in the oceanic front.

In Zone F (2.8–0 kyr BP), subsurface conditions were cool with stable, higher salinities as inferred by the dominance of C. reniforme and E. excavatum f. clavata, and stable abundances of N. labradorica, Buccella spp., E. bartletti, I. helenae, E. albiumbilicatum and I. norcrossi. N. labradorica and I. norcrossi indicate that although the LC water was the dominant water mass, Atlantic-derived water still entered the bay via the Slopewater Current. Productivity was high as evidenced by the high foraminiferal fluxes and in the high Ca and Sr concentrations in both cores 10G and 12G (Figure 4; Supplementary Figures S4–S5, available online), although local differences are seen between the core sites.

Discussion

Constructing a combined benthic foraminiferal record

This study is based on data from three marine sediment cores, which have been combined based on their chronologies to create a single composite record. Although the cores are from the same bay, they were not taken from the same exact site, so some limitations to this method should be considered. The XRF records from the three cores show a similar overall development where they overlap (Figure 4), thus strengthening the rationale of using the combined chronology. The assemblages and fluxes of benthic foraminifera match very well in general between the cores, although there are some intervals with very different fluxes; this could be due to preservation issues (e.g. core 10G may have lower foraminiferal preservation) as this core has the consistently lowest foraminiferal flux values. Differences in microhabitat distribution of benthic foraminifera (Corliss, 1985; Dijkstra et al., 2013) could also potentially cause differences in the foraminiferal assemblage from site to site. However, the overall foraminiferal records are very similar, and the general trends of the major species abundances are in good agreement, suggesting that this is not a significant issue and may at most have caused minor differences between sites. This has thus allowed us to combine the benthic foraminifera from the three cores into one single record of subsurface conditions (Figure 3).

For a detailed discussion of the YD interval, its two-part division (reflected in Zones A and B), and termination, we refer to Pearce et al. (2013, 2014a), who found that the first half of the YD was marked by cold surface conditions and a stratified water column, while the latter half was marked by variable sea-ice concentrations and northward migration of the LC–GS oceanic front. Together, these oceanographic changes predated the atmospheric changes used to mark the YD to Holocene transition.

Early Holocene warming

The high productivity and slightly reduced salinities in the early Holocene indicate that Placentia Bay was a biologically productive and hydrographically dynamic environment. The presence of N. labradorica indicates the presence of a nearby oceanic frontal zone facilitating high food availability and suggests that the oceanic front between the GS and the LC had pushed northward at the YD–Holocene transition, linked to a strengthening of the AMOC (Pearce et al., 2013). This northward movement of the front was presumably further accentuated at ~10.7 cal. kyr BP, when C. lobatulus indicated strengthened bottom-water currents. This may potentially be ascribed to a stronger Slopewater Current than today.

An increased GS influx and stronger AMOC are in agreement with a warm early Holocene as previously suggested among others for the north Atlantic (De Vernal and Hillaire-Marcel, 2006; Gibb et al., 2013), the Barents Sea (Duplessy et al., 2001) and the northern Baffin Bay (Knudsen et al., 2008), with sea-surface temperatures in the Baffin Bay Open Water Polynya at least as warm as today (Levac et al., 2001). An early Holocene strengthening of the AMOC is also in agreement with the high influx of relatively warmer subsurface waters to west and east Greenland (Jennings et al., 2006; Seidenkrantz et al., 2013).

Foraminiferal events – Unstable conditions due to meltwater pulses?

The early Holocene was a period of significant climatic variability. Until ca. 8 cal. kyr BP, the North Atlantic region experienced several cooling events interrupting the general warming trend. The best known climate fluctuation of that time is the so-called ‘8.2 BP event’, a short-lived period when major incidences of meltwater release and iceberg rafting from the melting ice sheets presumably slowed ocean circulation (AMOC), thereby causing a rapid and short-lived cooling of the northern hemisphere (Alley et al., 1997; Barber et al., 1999; Mayewski et al., 2004). However, cold events recorded in Greenland ice cores have also been described, centred around 9.95–9.3 kyr BP (Rasmussen et al., 2007).

Signal of Lake Agassiz drainage

A strong signal in the foraminiferal assemblages occurred from 8.4 to 8.25 cal. kyr BP, primarily characterized by a peak in G. auriculata arctica, a low-oxygen tolerating species also associated with productivity. The generally high foraminiferal abundances indicate that conditions were productive and calcareous foraminiferal tests were still well-preserved in the sediments. Results from the XRF analyses show a concurrent peak in Ti and Fe during the G. auriculata arctica peak (Figure 5), which may indicate increased input of terrestrial material. However, neither the grain size nor the dolomite and calcite records show much change during this event.

This event may be linked to the shift in oceanographic conditions that preceded the well-known 8.2 BP cooling event, which has been recorded around the North Atlantic (Anderson et al., 2007; Barber et al., 1999; Hillaire-Marcel et al., 2007; Keigwin et al., 2005; Klitgaard-Kristensen et al., 2001; Levac et al., 2011; Lewis et al., 2012; Young et al., 2012). It has been suggested that one or more catastrophic meltwater release events caused the cooling and that these meltwater events were derived from draining Laurentide glacial lakes (Barber et al., 1999; Hillaire-Marcel et al., 2007; Lewis et al., 2012; Roy et al., 2011). A clear freshwater influx signal, ascribed to a major drainage of Lake Agassiz at 8.4 kyr BP, was recorded in a core taken in the mouth of Hudson Strait (Barber et al., 1999). On the Laurentian Fan, a two-step cold pulse was recorded at ~8.5 kyr BP in alkenones and stable isotopes (Keigwin et al., 2005). This cooling and freshening of the sea surface lasted for ca. 700 years, overlapping our 8.4–8.25 kyr BP stratification event. Keigwin et al. (2005) linked this cold, relatively fresh pulse in surface water in the western north Atlantic to the outburst of glacial meltwater from Hudson Strait. This finding was corroborated by Levac et al. (2011) who identified two episodes of meltwater drainage at the 8.2 event based on dinoflagellate cyst investigations of sediment cores from the northeastern Newfoundland and northern Scotian shelves. Our 8.4–8.25 cal. kyr BP foraminiferal signal also supports a freshwater release, possibly from Hudson Bay–Lake Agassiz. This meltwater presumably travelled south with the LC along the western margin of the Labrador Sea across the Grand Banks (Fratantoni and McCartney, 2010; Lazier and Wright, 1993; Lewis et al., 2012). While the late YD H0 event can be traced in both XRF and XRD data as a detrital carbonate layer in Placentia Bay (Pearce et al., 2013, 2014a), the only geochemical evidence of the 8.4–8.25 kyr BP event is a minor peak in Ti and Fe (Figure 5), indicating that the meltwater pulse was either somewhat weaker or operated via a different mechanism than during the H0 event.

Unstable conditions from 9.7 to 9.2 kyr BP

The interval from 9.7 to 9.4 kyr BP is marked by the near-absence of benthic foraminifera (Figure 5), indicating unstable and inhospitable conditions in Placentia Bay. This occurred shortly before a peak in the low-oxygen tolerating species B. pseudopunctata at 9.2 kyr BP. Low bottom-water oxygenation is also indicated by the fact that most of the few foraminifera present in the low-abundance interval were etched, indicating increased acidity in the bottom water (Corliss and Honjo, 1981; De Vernal et al., 1992; Hald and Steinsund, 1996; Seidenkrantz et al., 2007).

Possible explanations for this event include slower currents (Kaiho et al., 1996; Murray, 1991), higher organic content at the seafloor (Murray, 1991), local riverine runoff, and/or a freshwater cap causing stratification of the water column. The largely unchanging grain size, here interpreted as a proxy for current strength (Gröger et al., 2003; Prins et al., 2002; Revel et al., 1996), and the relatively stable abundances of C. lobatulus during this time period do not indicate major changes in current regime. Ca and Sr values are high (Figures 4 and 5; Supplementary Figure S5, available online), while a low variability in dolomite and calcite (Figure 5) suggests no significant changes in mineralogical provenance in this part of the record. Ti and Fe show little variability in terrestrial input, indicating little change in river runoff (Figure 5). A mechanism that would explain the drop in foraminiferal abundance is strong stratification leading to reduced overturning of the water column and therefore less oxygen at the seafloor (causing few to no benthic foraminifera), followed by an increase in overturning (resulting in an increase in productivity and foraminifera). Therefore, we suggest that from 9.7 to 9.4 kyr BP, Placentia Bay was affected by strong water column stratification caused by a freshwater cap at the surface. This was followed by an increase in water column mixing around 9.2 kyr BP, which facilitated enhanced productivity with the opportunistic, low-oxygen tolerating species B. pseudopunctata among the first to colonize the area.

This freshwater cap might have been caused by a meltwater release coming from further north, along the path of the LC. Meltwater events recorded on the Labrador Shelf described by Jennings et al. (2015) could have caused the unstable conditions in Placentia Bay. In a core taken from Cartwright Saddle (Figure 1) on the Labrador Shelf, meltwater pulses from the Laurentide Ice Sheet were recorded in detrital-rich sediments beginning around 12.2 kyr BP (Jennings et al., 2015). Here, a well-defined detrital carbonate peak at 9.5 kyr BP (Jennings et al., 2015) suggests a major meltwater event likely originating in the Hudson Bay region. In fact, a whole suite of detrital carbonate events carried southward by the LC indicate the presence of more than one meltwater discharge event from Hudson Bay in the early Holocene (Jennings et al., 2015; Lewis et al., 2012; Roy et al., 2011). This is also supported by dinoflagellate cyst assemblages from the mouth of Hudson Bay that indicate overall unstable sea-surface salinities prior to 8.5 kyr BP, likely due to glacial discharge from the Laurentide Ice Sheet (De Vernal and Hillaire-Marcel, 2006). This was also seen as a cooling during the so-called ‘9.3 kyr event’ (9.31 b2k) that Rasmussen et al. (2007) identified in Greenland ice core records.

Such large-scale meltwater release events, with freshwater transported as a surface layer by the LC, resulted in freshening of the surface waters in the southwestern Labrador Sea, which may explain the strong water column stratification events in Placentia Bay. It is thus possible that the foraminiferal interval at 9.2–9.7 kyr in our record is also a distal signal of meltwater release from the Canadian Arctic (Anderson et al., 2007; Shaw et al., 2006). However, due to the large number of detrital carbonate events at this time (Jennings et al., 2015), it is not possible to link our foraminiferal event to one specific meltwater event with certainty. Nevertheless, as local Newfoundland glaciers had already disappeared by ca. 10 kyr BP (Dyke, 2004), a local source of freshwater is unlikely.

There is also a possibility that at least some of the meltwater could have originated in the Gulf of St. Lawrence, as it was highly variable during the early Holocene, due to meltwater runoff (Anderson et al., 2007; Levac et al., 2015); and at least during the YD, the Gulf of St. Lawrence outflow played a role (Levac et al., 2015). However, based on diatom assemblage investigations, Lapointe (2000) showed that any impact of meltwater from the Gulf of St. Lawrence was much reduced already by Cabot Strait. Furthermore, the surface waters in the Gulf of St. Lawrence showed a steady progression towards modern conditions with no major meltwater perturbations beginning ca. 9 kyr BP (10 ¹⁴C kyr BP) (Lapointe, 2000).

An oceanographical shift at 7.3 cal. kyr BP

After ca. 8 cal. kyr BP, the warm, saline GS waters had a relatively strong presence in the IC and WGC, creating productive environments around Iceland (Ólafsdóttir et al., 2010) and West Greenland (Lloyd et al., 2005; Seidenkrantz et al., 2013). Similar conditions may be inferred for the SW Labrador Sea based on our present study.

However, our study indicates that around 7.3 kyr BP (Subzone E1, Figure 3), an oceanographic regime shift occurred in the SW Labrador Sea. This was characterized by decreased bottom-water salinities and more unstable water column conditions as evidenced by an abrupt, significant increase of the abundance of E. bartletti in the beginning of Subzone E1. E. bartletti subsequently decrease slowly in abundance through the rest of Subzones E1–E2, while E. excavatum f. clavata and Buccella spp. gradually increase after ca. 6 kyr BP. At the same time, the bottom currents were weakened, as shown by the low abundances of C. lobatulus. This agrees well with the results of the dinoflagellate cyst study from core 12G (Solignac et al., 2011), which indicate relatively low surface-water salinities and possibly even the presence of sea ice carried by the LC during the mid-Holocene in Placentia Bay.

The colder and less saline surface-water conditions were attributed to enhanced transport of meltwater from the north (Levac et al., 2001; Solignac et al., 2011) because of increased melting of Arctic glaciers during the warm Northern Hemisphere summer climate of Holocene Thermal Optimum (HTO) (Solignac et al., 2011). We believe that this influx of cold, low salinity water also influenced subsurface conditions, inferring that our faunal change at 7.3 kyr BP illustrates the initiation of this enhanced meltwater transport from the north. This interpretation is supported by the relatively high values of Ti and Fe after 7.3 kyr BP (Figure 4; Supplementary Figures S4 and S5, available online), indicating an increased influx of terrestrial material into the LC. This increased strength of the LC, linked to increased melting of glaciers in the Canadian Arctic during the HTO, was originally suggested by Scott and Collins (1996); such increased glacier melting due to atmospheric warming is a well-known phenomenon (Alley et al., 2010; Zwally et al., 2002). Our benthic foraminifera (Figure 3) and elemental data (Figure 4) indicate that this influx of (possibly ice-loaded) meltwater continued, with gradually decreasing strength, until ca. 4.0 kyr BP (the start of E2), agreeing with Solignac et al. (2011).

Conditions in the western Labrador Sea contrasted those of the eastern margin, where meltwater release from the Greenland Ice Sheet decreased drastically around 7.5 kyr BP, or slightly thereafter (Ren et al., 2009; Seidenkrantz et al., 2013). This, however, was presumably largely linked to the dynamics of the retreating ice sheet rather than to a clear change in melting rate. Concurrently, the eastern Labrador Sea experienced an increase in bottom currents and productivity indicating a major reorganization of the Labrador Sea circulation around 7.5 kyr BP (Seidenkrantz et al., 2013). At the same time, or shortly after (~7 kyr BP), the overall circulation of the SPG changed, when modern circulation developed and deep water formation in the Labrador Sea commenced (Born and Levermann, 2010; Gibb et al., 2013; Hillaire-Marcel et al., 2001). A strong SPG would also facilitate a northward transport of warmer Atlantic-sourced water along the West Greenland coast and a strong southward flow of the LC off eastern Canada, likely pushing GS waters further off-shore. This shift at 7.3 kyr BP also involved the Norwegian Current branch of the North Atlantic warm water route in the Barents Sea (Berben et al., 2014), which implies a fundamental rearrangement of the North Atlantic heat and salt transport at that time.

The late Holocene

The foraminiferal faunas indicate that beginning around 4 kyr BP, conditions in Placentia Bay returned to warmer and more saline (Subzone E2). This indicates that although the LC remained the dominant water mass in Placentia Bay, the LC nevertheless weakened somewhat while influence from the Slopewater Current increased. The decreased southward flux of the LC would thus have allowed a northerly migration of the GC and the LC/GC frontal zone in this western part of the SPG. A decreased southward export of cold water by the LC after 4 kyr BP was also described from NE Newfoundland by Solignac et al. (2011) based on dinoflagellate cyst analyses. Solignac et al. (2011) primarily ascribed this weaker flow of cold water from the north to decreased melting of Arctic glaciers due to a general Northern Hemispheric cooling, thus resulting in less export of cold, ice-loaded water by the LC. A possible precursor to this change was described off Cape Hatteras, North Carolina, by Cléroux et al. (2012), who found increased surface-water salinities already by 5.2–3.5 kyr BP. Cléroux et al. attributed this to decreased export of colder waters from the north, allowing a northward shift of the GS.

The eastern side of the SPG, in contrast, experienced the opposite conditions: the East Greenland Current carried cold, low salinity surface waters through the Denmark Strait from 5 to 0.2 kyr BP (Jennings et al., 2011), with a reduction in the strength of the Irminger Current (IC) after 4.2 kyr BP on the SE Greenland shelf (Andresen et al., 2013). As the IC forms an important part of the North Atlantic circulation (Flatau et al., 2003), a weaker IC may also be linked to a general reduction in the AMOC. Although freshwater export via the Fram Strait strongly impacts Labrador Sea circulation and the LC (Cuny et al., 2002; Yashayaev, 2007; Zhang and Vallis 2006), the signal would be reduced before reaching Placentia Bay. This also supports our findings that the reduction in Arctic water and meltwater transport via the LC and subsequent warming at Newfoundland was not linked to a strengthening of the AMOC but rather a consequence of the general weakening of the AMOC and cooling of the North Atlantic region.

Subsurface conditions became more stable at ~2.8 kyr BP, as shown by increased C. reniforme and C. lobatulus becoming a consistent accessory species (Zone F). Surface waters were primarily affected by the LC but with some minor influence from warmer GS waters; these conditions were supported by a coeval increase in sea-surface salinities at ~2.8 kyr BP (Solignac et al., 2011). This shift at 2.8 kyr BP corresponds to the timing of the onset of the neoglacial cooling in SE and SW Greenland (Jennings et al., 2011; Møller et al., 2006), seen over large parts of the North Atlantic region (Koç et al., 1993), although some records show a cooling beginning as early as 3.5–3.6 kyr BP (Andresen et al., 2013) or 3.2–3.3 kyr BP (Funder and Fredskild, 1989; Koç et al., 1993).

The conditions in subsurface Placentia Bay were relatively stable during the last 2.8 kyr BP, showing only small changes in the foraminiferal assemblage composition, but with some variability in the benthic foraminiferal flux (Figure 3). The foraminiferal assemblages indicate a continued influx of LC water but with a still-important Atlantic water component. This suggests that the frontal zone between the GS and the LC was positioned near its present location, which may also explain the relatively productive environment. Jessen et al. (2011) linked a concurrent decrease in the abundance of exotic pollen from lower latitudes in Placentia Bay to a weakening of the southwesterly winds over the North Atlantic, implying a decrease in the advection of subtropical air masses after 3 kyr BP. They proposed a shift from a southwesterly wind–dominated atmospheric circulation pattern to one dominated by northern to northwesterly winds. At the same time, climate shifted to a generally more negative NAO scenario which would lead to reduced flux of Arctic water via the Canadian Arctic gateways. The overall decrease in the strength of the LC may thus be linked to the combined effect of a shift to a generally more negative NAO scenario, possibly combined with a further reduction of meltwater release from the Canadian Arctic, leading to warmer surface and subsurface temperatures along the path of the LC. The gradual decrease in sea-surface temperatures over the last 2000 years seen in the North Atlantic region (Cunningham et al., 2013; McGregor et al., 2015; Marcott et al., 2013; Pages 2k Consortium, 2013; Sicre et al., 2014; Wanamaker et al., 2008), possibly combined with an increase in LC strength (Sicre et al., 2014; Wanamaker et al., 2008), indicates a gradually weakening AMOC that, however, is not reflected in the benthic foraminiferal fauna in Placentia Bay.

Conclusion

Placentia Bay in southern Newfoundland has been strongly influenced by the LC and the GS over the last 13 cal. kyr. The early Holocene showed an initial warming, likely due to a northward shift of the frontal zone between the GS and the LC at the transition from the YD, linked to a strengthening of the AMOC. However, the benthic foraminiferal record indicates that increased stratification prevailed at ca. 9.7–9.2 and 8.4–8.25 kyr BP, resulting in low bottom-water oxygen levels. We suggest that these events were caused by glacial meltwater events releasing freshwater to the LC resulting in a low-density surface layer, which reduced convection.

Our data reveal a major oceanographic shift at 7.3 kyr BP, which appears to reflect a fundamental reorganization of the North Atlantic pole-ward warm-water transport pattern. A concurrent cooling and freshening of the surface waters off Newfoundland during the mid-Holocene is explained by an increased influence of the LC and increased transport of low-saline, cold meltwater, likely linked to increased meltwater release from glaciers in the Canadian Arctic during the Northern Hemisphere HTO.

As the overall Northern Hemisphere climate cooled at the end of the HTO, the salinity of Placentia Bay bottom waters increased again and the LC weakened slightly causing a minor northward movement of the oceanic polar front between the LC and the GS. The late Holocene thus saw a minor northward movement and increased local influence of the GS, also contributing to a more saline and productive environment in Placentia Bay. However, this shift was not linked to a general strengthening of the AMOC but rather to the post-HTO Northern Hemisphere cooling, causing reduced transport of meltwater. The LC was again strengthened during the last ~3 kyr, creating the cold and low-saline conditions that persist today. A general shift in wind direction linked to an overall NAO-type climate leading to predominantly northerly wind may also have facilitated a stronger LC.

Footnotes

Acknowledgements

We thank the captain and crew of RV ‘Akademik Ioffe’ as well as the entire scientific party for their help during the research cruise. We also thank Anne Hansen Thoisen, Aarhus University, for help with x-ray fluorescence core scanning as well as x-ray diffraction (XRD) and grain size measurements. Thanks also to Njáll F Reynisson for help with the preliminary age-depth models, Elsa Picón Pineda for carrying out a foraminiferal pilot study of core 10G, Jesper Olsen for the age model of core 12G and Ole Bjørslev Nielsen for his help interpreting the XRD results.

Funding

This study was funded by ‘Kommisionen for Videnskabelige Undersøgelser i Grønland’ and by the Danish Council for Independent Research, Natural Science (project nos. 09-069833/FNU, 12-126709/FNU DFF-4002-00098_FNU). The cruise was funded by the Danish Council for Independent Research, Natural Science (project no. 272-06-0604/FNU) and carried out within the Danish–Russian collaboration project ‘Joint paleoceanographic investigations in the Labrador Sea region’. The research leading to these results has also received funding from the European Union’s Seventh Framework programme (FP7/2007-2013) under grant agreement no. 243908, ‘Past4Future, Climate change – Learning from the past climate’.

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