Site U1308 is located 2800 km from the mouth of Hudson Strait and represents a distal location for deposition by icebergs originating from the Labrador Sea. Heinrich layers are relatively thin in distal deposits, and consequently, bioturbation can easily blur the signal or obliterate internal structure of the event. To assess stratigraphic integrity and resolution, we used X‐ray computed tomography (X‐ray CT) to conduct a detailed analysis of sediment structure and bioturbation by examining 1 mm slices of cores from multiple holes at Site U1308. Core scanning XRF was conducted at 1 mm in Holes U1308A, U1308B, and U1308C, and discrete samples at 5 mm intervals were taken from Hole U1308A for IRD and foraminiferal point counting, faunal assemblages, mineralogy (X‐ray diffraction, XRD), and stable isotope analysis.

Location of Site U1308 in the IRD belt (blue shade) relative to reconstructed ice sheet extent in the Northern Hemisphere (continuous solid black line) showing the inferred locations of ice streams (black arrows; after]) at the Last Glacial Maximum. HS = Hudson Strait; CS = Cumberland Sound; LC = Laurentian Channel; and FM = Fleuve Manche. Figure modified after]. Also shown is the Iberian Margin (IM) and location of Site 976.

Here we examined the detailed structure of Heinrich Stadial 1 at International Ocean Discovery Program (IODP) Site U1308 (49°52.6661′N; 24°14.2875′W) (Figure 1 ), which is located within the ice‐rafted detritus (IRD) belt of the North Atlantic [ Ruddiman , 1977 ]. Site U1308 represents the reoccupation of DSDP Site 609, which has played an important role in the recognition of Heinrich events and correlation of millennial‐scale climate variability between marine sediment and Greenland ice cores [ Broecker et al ., 1992 ; Bond et al ., 1992 , 1993 , 1999 ; McManus et al ., 1994 ; Bond and Lotti , 1995 ].

In distal deposits in the open North Atlantic, Heinrich layers consist mainly of IRD and are almost devoid of planktonic foraminifera [ Heinrich , 1988 ; Broecker et al ., 1992 ]. Rashid et al . [ 2003 ] suggested that distal events are recorded as a single IRD layer that corresponds to only part of the sequence found in cores proximal to Hudson Strait. However, some detailed studies of distal HEs have suggested that the events may be more complex than a single episode of ice rafting. For example, a record of IRD from the Iberian Margin shows two distinct peaks in IRD during HE1 (i.e., H1a and H1b) [ Bard et al ., 2000 ], although only the older of these peaks contains DC (H1b). The younger H1a contains hematite‐coated grains (3%) but almost no detrital carbonate. On the Laurentian Fan, Gil et al . [ 2015 ] identified two ice‐rafting events during H1 that are associated with strong cooling and expansion of sea ice. Similar to the Iberian Margin, the older peak is rich in detrital carbonate and the younger is rich in hematite‐stained grains.

Although Heinrich layers were first identified in the open North Atlantic [ Heinrich , 1988 ], they have been correlated with detrital carbonate deposits from the Labrador Sea [ Andrews and Tedesco , 1992 ; Stoner et al ., 1996 ]. Deposits close to Hudson Strait show a more complete history of Heinrich layers but are affected by depositional processes other than ice rafting [ Hesse et al ., 2004 ; Rashid et al ., 2012 ; Hesse , 2016 ]. The complexity of the deposits is related to the release of subglacial meltwater that entrains sediment, triggering the transport of fine material within nepheloid flows and turbidity currents [ Hesse and Khodabakhsh , 1998 , 2016 ; Rashid et al ., 2003 , 2012 ]. Such flows provide an additional mechanism for the delivery of detrital to the deep sea, resulting in several different types of deposits in the Labrador Sea [ Hesse , 2016 ].

Heinrich layers (HLs) in North Atlantic sediments are formed from massive discharges of the Hudson Strait ice stream that drained the Laurentide Ice Sheet [ Heinrich , 1988 ; Bond et al ., 1992 ; Broecker et al ., 1992 ; Hemming , 2004 ]. These layers are identified in North Atlantic sediment cores by poorly sorted sediments that are almost devoid of foraminifera and contain large amounts of ice‐rafted detritus (IRD), which is rich in detrital carbonate (DC) sourced from Paleozoic limestone and dolostone from the region of Hudson Bay and Hudson Strait [ Broecker et al ., 1992 ]. The layers are found from the Labrador Sea in the western Atlantic to off the coast of Portugal in the eastern North Atlantic along a zonal axis between 40 and 50°N in the so‐called “Ruddiman IRD belt” [ Ruddiman , 1977 ]. The deposits thin eastward from 40–70 cm in the Labrador Sea to <5 cm in the central North Atlantic, to <1 cm in the east [ Dowdeswell et al ., 1995 ].

The literature is confounded by different usages of the terms Heinrich layer (HL), Heinrich event (HE), and Heinrich stadial (HS). Following suggestions from other authors [ Barker et al ., 2009 ; Sanchez Goñi and Harrison , 2010 ], we adopt consistent definitions and usage of these terms throughout this paper as described below. A Heinrich event (sensu stricto) is the climate event produced by a massive discharge of the Laurentide Ice Sheet through Hudson Strait. Heinrich events occurred during some but not all stadials, which were part of millennial‐scale Dansgaard‐Oeschger (D‐O) oscillations during the last glaciation. A Heinrich layer is the physical manifestation of a Heinrich event preserved in a marine sediment core. A Heinrich stadial is the cold period that contains the Heinrich event. For the last deglaciation, Heinrich Stadial 1 is synonymous with the “Mystery Interval” (17.5 to 14.5 ka) [ Broecker and Barker , 2007 ]. A Heinrich event and Heinrich stadial are not the same because Heinrich events are shorter in duration than the Heinrich stadials in which they occur.

The peak of HL1.1 is estimated at 16.2 ka with a 2 sigma range from 15.5 to 17.1 ka. The age of ~16.2 ka for H1.1 agrees well with estimates of 16.13 ± 0.12 ka from an ice core drilled into the West Antarctic Ice Sheet (WAIS) [ Rhodes et al ., 2015 ], 16.2 ka from the Cariaco Basin off Venezuela [ Deplazes et al ., 2013 ], and 16.07 ka from Hulu Cave, China [ Treble et al ., 2007 ; Southon et al ., 2012 ; Zhang et al ., 2014 ]. The peak of HL1.2 is estimated at 15.1 ka with a range of 14.3 to 15.9 ka. The age model gives an age of 14.7 ka for the base of the Bolling‐Allerod warming in the core that matches the date obtained in the Greenland ice core chronology [ Svensson et al ., 2008 ].

We produced a radiocarbon age model by combining dates on N . pachyderma from Site U1308 with similar analyses from Site 609 [ Bond et al ., 1992 ]. The two sites were correlated precisely to one another using color reflectance [ Hodell et al ., 2008 ; Obrochta et al ., 2012 , 2014 ]. We used estimates of high‐latitude North Atlantic reservoir ages and errors given by Stern and Lisiecki [ 2013 ]. The age model was developed with Bacon [ Blaauw and Christen , 2011 ], which uses Bayesian statistics to reconstruct the accumulation history of the profile. Errors in both radiocarbon calibration and estimation of reservoir age are considered. Results include the mean, median, maximum, and minimum ages for each level in the profile (see supporting information ). The age model likely overestimates the duration of HL1.1 and 1.2, which were deposited rapidly and likely lasted for only a few hundred years [ Francois and Bacon , 1994 ; Thomson et al ., 1995 ; McManus et al ., 1998 ].

The radiocarbon dates from Hole U1308A (Table 1 ) agree well with radiocarbon ages reported by Bond et al . [ 1992 ] for Site 609 (see supporting information ). At Site 609, the two detrital carbonate peaks occur at 81.5 and 78.5 cm and the nearest radiocarbon dates are at 84–85 cm (14,990 ± 230) and 79–81 cm (13,650 ± 90). The radiocarbon dates from Site 609 are in close agreement with the basal dates for H1.1 (14,762 ± 45) and H1.2 (13,355 ± 42) obtained for Hole U1308A.

The younger DC layer (H1.2) occurs at 81.5 cm with a radiocarbon date of 13,355 ± 42 Δ 14 C years B.P. at 82.5–83 cm just below its base (Figure 7 ). The calibrated age is between 15.6 and 15.2 ka assuming a reservoir age of 500 years for this time period [ Stern and Lisiecki , 2013 ] (Table 1 ). HL1.2 has a sharp base but shows evidence of bioturbation at the top of the layer (Figure 11 ). Radiocarbon ages at 81.75 cm from the peak of HL1.2 are 13,655 ± 30 14 C years B.P. for N . pachyderma and 12,815 ± 30 14 C years B.P. for G . bulloides . The calibrated age range for N . pachyderma is 16 to 15.6 ka assuming a 500 year reservoir age and 15.2 to 14.8 ka assuming a 1000 year reservoir age. Radiocarbon dates for G . bulloides suggest a younger age (13.9 to 13.7 ka) for HL1.2 that is within the Bolling‐Allerod. N . pachyderma (sin) and G . bulloides may have different reservoir ages given that the environmental tolerances and depth habitats differ between the two species. Alternatively, specimens of G . bulloides may have been mixed down into HL1.2 from the Bolling‐Allerod interval above [ Manighetti et al ., 1995 ]. The similar ages of both foraminiferal species between 86 and 81 cm may reflect rapid deposition between HL1.1 and 1.2 or a radiocarbon plateau identified by Sarnthein et al . [ 2015 ] between 16.05 and 15.25 ka (Figure 7 ). The plateau presents difficulty in accurately dating HL1.2.

The lower DC layer (H1.1) has a sharp base indicating little mixing of material from below the IRD layer (Figures 2 and 3 ). The uncorrected radiocarbon age obtained just below the increase in IRD is 14,762 ± 45 Δ 14 C years B.P. at 88.5–89 cm (Figure 7 ). The calendar age is dependent upon the assumption made for the reservoir age correction [ Reimer et al ., 2013 ]. It is likely that reservoir ages in the glacial North Atlantic were significantly greater than the modern value of 400 years during Heinrich Stadial 1. Stern and Lisiecki [ 2013 ] suggested that the reservoir age in the North Atlantic during the early deglacial increased to >1000 14 C years between 18.5 and 16.5 kyr B.P. Sarnthein et al . [ 2015 ] suggested a reservoir age of 1300 to 1650 14 C years during HS‐1 in the subtropical North Atlantic. Assuming a reservoir age of 1200 years [ Stern and Lisiecki , 2013 ], the calibrated age at 88.5–89 cm is between 16.2 and 16.5 ka (2 sigma), suggesting that the onset of H1.1 is younger than this age range. This represents a maximum age because bioturbation is suppressed by rapid deposition of the Heinrich layer [ Manighetti et al ., 1995 ]. The next dated level at 86.5 to 87 cm occurs where the %IRD is high but N . pachyderma (s) has decreased to 60%. The radiocarbon age of 13,915 ± 30 14 C years gives an age range of 15.6 to 15.2 (2 sigma) assuming a 1000 year reservoir age. This represents a minimum age because Heinrich layers are virtually devoid of foraminifera and downward bioturbation of foraminifera from above is likely (Figure 8 b) [ Manighetti et al ., 1995 ].

Radiocarbon dating was performed in Section U1308A‐1H‐1 across HL1.1 and HL1.2 to constrain the ages of the two IRD layers. The polar form N . pachyderma was measured in eight samples, and G . bulloides was analyzed in four samples. The ages of N . pachyderma are consistently older than G . bulloides at the same stratigraphic level by an average of 650 years (Figure 7 and Table 1 ). Radiocarbon ages decrease from 92 to 86 cm and remain relatively constant from 86 to 80 cm before decreasing again to 78 cm (Figure 7 ).

Our bioturbation analysis suggests that the distinct double peak in Hole U1308A was primary in origin and not caused by upward dispersal of detrital carbonate from the lower peak by bioturbation. Moreover, the absence of a double peak in Hole U1308B is attributed to bioturbational disturbance that has mixed the two peaks. We therefore focused our efforts on Hole U1308A because it has been less affected by bioturbation than Hole U1308B.

Trace fossil distribution and bioturbation index (BI) in the studied Holes U1308A and U1308B, through the differentiated intervals 18 to 7. Note the absence of discrete trace fossils in intervals 12 and 13 (BI = 0) of Hole U1308A and the presence of Thalassinoides and Planolites (BI = 2–3) in the corresponding intervals in Hole U1308B.

Bioturbation was studied by digital image analysis [ Dorador and Rodríguez‐Tovar , 2014 ; Dorador et al ., 2014a , 2014b ; Rodríguez‐Tovar and Dorador , 2015 ] in Holes U1308A and U1308B to determine why Hole U1308A records a distinct double peak, whereas U1308B has a broad single peak. Hole U1308B has a greater overall abundance of large traces fossils and greater bioturbation index than Hole U1308A (Figure 11 ). Thalassinoides is more abundant in Hole U1308B than it is in U1308A, whereas Planolites is more dominant at Hole U1308A than U1308B. Hole U1308A contains some intervals with a bioturbation index of 0 (bioturbation absent) that is not observed at Site U1308B. Importantly, the interval between HL1.1 and HL1.2 (intervals 12 and 13 in Figure 11 ) in Hole U1308A is lacking trace fossils, whereas the corresponding interval in Hole U1308B contains large traces of Thalassinoides (Figure 11 ).

Benthic δ 18 O also decreases across the base of HL1 marking the start of deglaciation (Figure 10 ). At Hole U1308A, the detrital carbonate peaks are marked by two minima in benthic δ 18 O. Benthic δ 13 C decreases prior to HL1.1 and remains low through HL1.1 and HL1.2 before increasing at the start of the Bølling‐Allerød.

The δ 18 O of N . pachyderma (sin) decreases below the base of HL1 beginning ~90 cm and reaches minimum values at 87.25 cm in Hole U1308A (Figures 9 and 10 ). Oxygen isotope values increase slightly in the interval between HL1.1 and HL1.2. The δ 18 O of Hole U1308A is very similar to the records from Holes U1308C and 609 (Figure 10 ).

The δ 18 O of bulk carbonate was measured using the method of Spötl and Vennemann [ 2003 ] and reflects the relative proportion of detrital and biogenic carbonate [ Hodell and Curtis , 2008 ] and shows two distinct lows during HL1, reaching minimum values at 87.5 and 82.5 cm (Figure 9 ). Bulk carbonate oxygen isotopes have negative values reaching as low as −5‰. These values are outside the range of biogenic carbonate seen at Site U1308 and are diagnostic of detrital carbonate derived from Hudson Strait [ Hodell and Curtis , 2008 ].

Planktonic foraminifera were counted using standard methods [ Vautravers et al ., 2004 ; Vautravers and Shackleton , 2006 ] following the taxonomy of Kennett and Srinivasan [ 1983 ]. Foraminiferal assemblages are nearly 100% N . pachyderma (sin) below HL1 (Figure 8 ). The abundance of foraminifera is at a minimum, and % N . pachyderma (sin) remains near 100% during HL1.1. Toward the top of HL1.1 at 86.5 cm, the percentage of N . pachyderma (sin) declines abruptly as the temperate planktonic foraminifera increase, indicating a rapid warming of surface waters. The sediment between the two IRD peaks of H1 contains abundant foraminifera consisting of a temperate water assemblage with relatively low percentages of N . pachyderma (sin), averaging 30% of the total assemblage. A small increase in N . pachyderma (sin) and decrease in foraminiferal abundance is associated with the younger IRD peak (H1.1) at 82.0 cm.

A double peak also occurs in detrital carbonate and IRD during H1 at Site 609 [ Bond et al ., 1992 , 1993 ; Bond and Lotti , 1995 ], but the sampling resolution at Site 609 was 1 cm rather than 0.5 cm used at Site U1308. Nonetheless, two peaks in % detrital carbonate at 78.5 and 81.5 cm at Site 609 match the peaks at Site U1308 (see supporting information ). The lower IRD layer is the thicker of the two and has a greater concentration of IRD grains and dolomite.

%IRD in the >150 µm size fraction determined by point counting in Section U1308A‐1H‐1. %Dolomite and Quartz were determined by XRD of bulk sediment (). The peak in %quartz at 89–90 cm is a precursor to Heinrich Events 1.1 and 1.2 that are marked by the peaks in %IRD and dolomite.

Point counting of the >150 µm size fraction in Hole U1308A confirmed the two peaks in DC (Figures 6 and 7 ), corresponding to the IRD layers that are observed in the CT images and recorded by elemental ratios (Ca/Sr) measured using the XRF scanner (Figure 5 ). In Hole U1308A, % total IRD counts begin to increase between 88 and 88.5 cm. A peak in quartz at 89–89.5 cm precedes the increase in detrital carbonate at 87–88 cm (Figure 6 ) and corresponds with a peak in Si/Sr (Figure 5 ). This early IRD is rich in quartz and may represent a “precursor event” as observed at other sites in the North Atlantic [ Grousset et al ., 2001 ].

Mn/Al shows three peaks in the top 150 cm of Site U1308. Two peaks occur between Heinrich Layers 1.1 and 2 with one peak above H2 and the other near the base of HL1.1 (Figure 5 ). The third peak occurs at about 65–70 cm above HL1.2.

Si/Sr reflects the relative contribution of silicate‐rich IRD (Si) and biogenic carbonate (Sr) and has been shown to correlate well with %lithics or total lithics per gram at Sites 609 and U1308 [ Hodell et al ., 2008 ; Obrochta et al ., 2014 ]. High Si/Sr values indicate increased abundance of silicate‐rich IRD at Site U1308 [ Hodell et al ., 2008 ]. Si/Sr is high within both DC‐rich layers of Heinrich 1 and is also elevated in the interval leading up to H1 (Figure 5 ). There is an increase in Si/Sr at 97 cm in Hole U1308B and 107 cm in Hole U1308A during HS1, which define the end of the Last Glacial Maximum and beginning of Heinrich Stadial 1.

Hodell et al . [ 2008 ] demonstrated that Ca/Sr is a reliable proxy for the concentration of detrital carbonate at Site U1308. HL1 is marked by increases in Ca/Sr in all three holes, but subtle differences exist in the structure of the signal among holes (Figure 5 ). In Hole U1308A, HL1 consists of two sharp peaks in Ca/Sr corresponding with the density peaks in the CT images. The lower, older (H1.1) peak is the stronger of the two and is about 2 cm wide, whereas the upper, younger (H1.2) peak is 1 cm wide. The two peaks are separated by about 3 cm of intervening sediment that contains no detrital carbonate. In Hole U1308B, HL1 consists of a single diffuse Ca/Sr peak that is spread over 7 cm. In Hole U1308C, HL1 consists of two peaks including a larger older peak and smaller second peak separated by less than a centimeter of sediment. The second peak is much weaker and less well defined in Hole U1308C than it is in U1308A. There is also some indication that HL 2 may be double peaked (Figure 5 ), but we have not studied this event in any detail.

The quality of XRF data is highly dependent on the physical properties of the sediment core (e.g., varying porosity) [ Grützner and Higgins , 2010 ; Röhl and Abrams , 2000 ] and condition of the split core sediment surface and is affected by mineral heterogeneities and surface roughness [ Tjallingii et al ., 2007 ]. This is particularly problematic for sediment containing IRD because a single grain at the surface can cause large anomalies in elemental variations. Comparison of XRF scans in multiple holes is necessary to evaluate the reproducibility and integrity of the XRF results.

The number of grains of IRD in the CT images was automatically counted using FIJI (ImageJ) software [ Schindelin et al ., 2012 ] by thresholding each 1 mm slice to exclude fine‐grained, low‐density matrix sediment and isolate individual IRD grains (Figure 4 ). IRD grains (>1 mm in size) that were large enough to cross images consisting of 1 mm slices were counted. Grains were counted using a plug‐in for FIJI called BoneJ [ Doube et al ., 2010 ], which hosts an object counter called Particle Analyzer. The plug‐in is extremely efficient at quantifying the particles in an image stack but only recognizes particles that span more than one image (i.e., >1 mm in size). IRD counts were binned in 5 mm intervals so that they could be compared with actual counts of IRD obtained from the >150 µm fraction in discrete samples taken every 5 mm (see section 3.4 ). The pattern of the automated grain counts of >1 mm shows two distinct peaks as do the IRD counts from the >150 µm fraction (Figure 4 ).

(a) Comparison of IRD point counts from the >150 µm fraction (black) and IRD > 1 mm estimated by image analysis of the X‐ray CT images (red). (b) Density estimated by CT core scanner (filled gray area) and gamma ray attenuation (blue) for Section U1308A‐1H‐1 between 70 and 100 cm. The sediment thicknesses at half peak of H1.1 and 1.2 and the intervening interval are indicated.

Density was calculated for each 1 mm X‐ray CT slice in Hole U1308 by averaging the gray scale of all pixels after removing holes and cracks from each slice (Figure 4 ; see supporting information ). In Hole U1308B, the lower layer is up to 2 cm thick, whereas the upper layer is 1 cm thick and the top is bioturbated. The layers also correspond to increases in wet bulk density of the core measured during Expedition 303 by gamma ray attenuation (GRA), which reaches 1.8 g cm −3 in each of the two peaks of H1 in Hole U1308A (Figure 4 ) [ Expedition 303 Scientists , 2006 ]. The GRA was measured at 2.5 cm intervals in Hole U1308A, and the peaks are defined by single points, and fine‐scale structure is not captured.

X‐ray CT images were acquired for Holes U1308A and U1308B (see supporting information ) but could not be obtained for U1308C because of heavy sampling of both archive and working halves of the core. The CT images reflect density contrasts such that ice‐rafted detritus has a greater density than the background pelagic sediments that consists mostly of nannofossil ooze and clay [ Expedition 303 Scientists , 2006 ]. Two dense layers are distinctly visible in the CT scan of Hole U1308A, whereas a single mixed layer occurs in Hole U1308B (Figure 3 and supporting information ). We refer to the lower peak as H1.1 and the upper peak as H1.2 to distinguish them from H1a and H1b defined by Bard et al . [ 2000 ] because we are not certain they represent the same events. Each Heinrich layer is poorly sorted, consisting of coarse‐grained clasts of IRD embedded in a dense fine‐grained matrix. Some large millimeter‐sized clasts of IRD are visible in the dark layers immediately below the base of the Heinrich layer in both holes.

Heinrich Layer 1 was recovered in IODP Holes U1308A, U1308B, and U1308C and in two holes at DSDP Site 609. All core images show similar features including a sharp color contrast at the base (Figure 2 and supporting information ), representing the contact between light colored carbonate‐rich sediment of the Heinrich layer and a dark layer below that is rich in manganese. Heinrich Layer 1 is overlain by lighter colored sediments deposited during the Bølling‐Allerød (B‐A) followed by a transition to the Younger Dryas (Y‐D).

4 Discussion

4.1 IRD in a Dense Fine‐Grained Matrix In Hole U1308A, Heinrich Event 1 consists of two discrete density peaks (Figure 4). The coarse IRD grains of H1.1 and H1.2 are set within a dense matrix of fine‐grained sediment, which is expressed in the CT slices by a cloud of high‐density material (Figure 3; also see supporting information). We interpret the fine sediment as glacial flour that is released and dispersed from melting icebergs along with coarser‐grained material [Gilbert, 1990]. Glacial flour or rock flour is clay‐ and silt‐sized sediment produced by mechanical grinding of bedrock by glaciers. Glacial‐aged icebergs derived from the Laurentide Ice Sheet likely had a greater proportion of fine‐grained carbonate than modern Atlantic icebergs, which are derived mainly from the Greenland Ice Sheet, because of efficient grinding of Paleozoic carbonates that underlie Hudson Strait and Hudson Bay [Hiscott and Aksu, 1996; Hesse and Khodabakhsh, 2016; Hesse, 2016]. Although IRD layers are often identified through properties of the coarse‐sediment fraction, the sand fraction is usually a relatively small portion of most glacial‐derived sediment, which is typically in the silt and clay size range [Andrews, 2000; Andrews and Principato, 2002]. This smaller size fraction is missed in counts of coarse‐grained IRD but is a major contributor to glacimarine sedimentation. Glacial flour and especially carbonate can remain in suspension for a long time because of its fine grain size and resistance to flocculation. It can be transported to great distances by currents after release from icebergs. As a result, glacial flour was likely widely dispersed throughout the North Atlantic IRD belt during Heinrich events. The glacial flour increases water turbidity during Heinrich events and, together with low temperatures and increased stratification, suppresses primary productivity that would further hinder this fine material from being incorporated into fecal pellets and sinking [Syvitski, 1989]. As a result, the wide distribution and low δ18O values of the detrital carbonate from Hudson Strait are particularly useful for detecting Heinrich events within and outside the margins of the North Atlantic IRD belt [Hodell and Curtis, 2008]. Within the poorly sorted IRD layers, the fine‐grained glacial flour fills the voids between larger particles resulting in a substantial increase in sediment packing, thereby resulting in high bulk density (Figure 4). The high density may also be a result of partial cementation of Heinrich layers by authigenic dolomite and ankerite [Tamburini et al., 2002]. Heinrich layers were originally described as “cemented marls” in German cruise reports [Heinrich, 1988; Hemming, 2004], and scanning electron microscope photos of Heinrich layers at Site U1308 reveal rhombohedral crystals of dolomite, suggesting that at least some of the dolomite is authigenic (see supporting information).

4.2 Twin Peaks Our results from Hole U1308A indicate that Heinrich Layer 1 consisted of two discrete pulses of IRD (Figure 6). The two peaks are not the result of bioturbation or coring disturbance. The twin peaks are also observed in the low‐resolution counts of detrital carbonate of Bond et al. [1992] from nearby Site 609, supporting the dual nature of H1 (see supporting information). Other records also exhibit twin peaks for H1 across the North Atlantic indicating that it may be a pervasive feature. For example, Heinrich layers typically feature two peaks in coarse‐grained detrital carbonate within the Labrador Sea [Rashid et al., 2003]. This pattern has been attributed to the dilution of the middle of a single IRD layer by meltwater input of fine material [Hesse and Khodabakhsh, 1998, 2016; Clarke et al., 1999; Rashid et al., 2003]. However, a record from sediment core GGC31 on the crest of Orphan Knoll (water depth 1796 m) also shows the pattern of two peaks in detrital carbonate as observed at Site U1308 [Bond and Lotti, 1995]. This core is relatively shallow, and located on an isolated high, and so is not affected by turbidity currents originating from Hudson Strait and is too far from Hudson Strait to be affected by meltwater plumes derived from the ice sheet terminus [Hesse, 1989; Hesse and Khodabakhsh, 2016]. Similarly, Core EW9302‐2JPC (1251 m, 48°47.70′N, 45°05.09′W) on the Flemish Cap also shows two distinct peaks during H1 [Marcott et al., 2011]. Site U1302/U1303 at 3600 m water depth on Orphan Knoll also shows a small secondary peak associated with the main H1 event (see supporting information). Sediment core SU90‐09, on the southern fringe of the IRD belt, displays two peaks in the proportion of IRD that is detrital carbonate within H1 [Grousset et al., 2001]. A record from core SU81‐18 from the Iberian Margin shows two distinct IRD peaks that coincide with those seen at Site U1308, although only the earlier of the two peaks is rich in detrital carbonate [Bard et al., 2000]. Bard et al. [2000] provided ages of 16.0 (H1a) and 17.5 ka (H1b) for the two peaks, assuming a reservoir age of 400 years for Core SU81‐18 on the Iberian Margin. However, Skinner et al. [2014] demonstrated that reservoir ages on the Iberian Margin were not a constant 400 years but varied by up to 1800 years during Heinrich Stadial 1. Applying a reservoir correction of >1000 years to Core SU81‐18 would make H1b significantly younger than 17.5 ka. We suggest that H1b of Bard et al. [2000] is synchronous with H1.1 at Site U1308. It is the stronger of the two peaks and likely occurred at 16.1 ± 0.123 ka [Rhodes et al., 2015]. It is uncertain whether H1a of Bard et al. [2000] corresponds to H1.2 because it does not contain detrital carbonate on the Iberian Margin. Two peaks in IRD are similarly recorded on the Laurentian Fan [Gil et al., 2015] and dated at 17.5 and 16 cal kyr B.P. These dates are also based on an age model that used a standard marine reservoir correction of 400 years [Keigwin et al., 2005].

4.3 Implication of Twin Peaks for Ice Sheet Dynamics The two DC‐rich IRD layers during H1 in Hole U1308A suggests that rapid iceberg discharge from the Hudson Strait ice stream may have occurred in two stages or, alternatively, it involved two independent ice streams. Alley and MacAyeal [1994] predicted that all Heinrich events should have a double‐peak structure related to the process of freeze on and incorporation of debris into glacial ice. According to their “binge‐purge” model [MacAyeal, 1993], ice calved early during the surge is debris laden from the previous freeze‐on portion of the binge‐purge cycle. When debris‐laden basal ice is completely melted off the ice stream, clean icebergs are produced resulting in continued freshwater forcing but a hiatus occurs in IRD deposition, producing a “midpurge hiatus.” After freeze on commences but before the purge ends, IRD flux resumes and continues to the end of the purge. Thus, three stages were proposed by Alley and MacAyeal [1994]: (i) the initial IRD output, (ii) a midpurge hiatus devoid of IRD, and (iii) a late purge IRD output. The IRD hiatus occurs because debris‐laden basal ice is completely melted from the base of the ice stream during the middle portion of the purge. Although the two peaks recorded at Hole U1308A appear to support the prediction of Alley and MacAyeal [1994] that IRD flux should reach a maximum twice during each Heinrich event, there are aspects that may not fit the model. According to the model, IRD is absent during the midpurge hiatus but meltwater is still produced from the melting of clean bergs. Planktonic foraminifera assemblages in the intervening sediment between H1.1 and H1.2 do not support the presence of cold meltwater at Site U1308 but instead are dominated by subpolar species, especially G. bulloides (Figure 8). The warming between H1.1 and H1.2 at Site U1308 does not preclude the possibility of continued iceberg delivery and cooling in the western Atlantic closer to Hudson Strait, but not in the central North Atlantic. Another consideration is the time difference between the two peaks. Alley and MacAyeal [1994] estimated the total duration of the Heinrich event to be 750 years with the debris hiatus lasting 250 years. The age difference between the base of H1.1 and H1.2 is 1400 radiocarbon years, but the absolute age difference between the two peaks depends upon assumptions about the reservoir correction for each date. If the two events occurred at ~16.2 and 15.1 ka, then 1100 years is longer than the predicted duration of 250 years for the midpurge hiatus. The problem is that the 95% age ranges for HL1.1 (15.5–17.1 ka) and HL1.2 (14.3 to 15.9 ka) overlap. If H1.2 is toward the older age of 15.9 ka, then the duration of the midpurge hiatus would be consistent with the estimate of Alley and MacAyeal [1994]. Alternatively, the two peaks may represent rapid iceberg delivery from different ice streams, both of which would be underlain by bedrock containing detrital carbonate. The older and main event involved the Hudson Strait Ice Stream. The younger event could be linked to sources from any number of other ice streams draining the eastern margin of the Laurentide Ice Sheet (Figure 1); for example, Rashid et al. [2012] suggested that discharge from the Hudson Strait Ice Stream in H1 was followed by the Cumberland Sound Ice Stream (Figure 1). A potential problem with this suggestion is that ice sheet basins such as those draining Cumberland Sound are 1 to 2 orders of magnitude smaller that the Hudson Bay‐Hudson Strait system, making the delivery of very large numbers of icebergs unlikely. Another possibility is the Laurentian Channell Ice Stream (Figure 1), which has a relatively large drainage‐basin area [Dowdeswell et al., 1995]. However, although the St. Lawrence valley contains limestone, we would also expect other lithologies (e.g., Permian‐Carboniferous red beds).

4.4 Warming During HE1 Several studies have reported warming during or preceding Heinrich events either in the surface or subsurface waters [Jonkers et al., 2010; Marcott et al., 2011; Naafs et al., 2013]. Warming in the subsurface has been observed in several models during Heinrich events and has led to the suggestion that subsurface warming may have been responsible for destabilization of grounded ice shelves followed by surging of ice streams [Hulbe, 1997; Hulbe et al., 2004; Álvarez‐Solas et al., 2010; Marcott et al., 2011; Bassis et al., 2017]. A possible problem for models invoking the collapse of extensive floating ice shelves in the delivery of IRD is that strong basal melting, which is common close to ice shelf grounding lines [e.g., Enderlin and Howat, 2013; Rignot et al., 2013], means that icebergs derived from the floating ice shelves themselves may be largely devoid of basal debris [e.g., Alley et al., 2005]. Debris‐rich icebergs with significant thicknesses of basal debris more probably come from icebergs sourced from grounded ice streams. Alternatively, ice sheet models can simulate a Heinrich event in response to subsurface warming without the necessity of a floating ice shelf [Bassis et al., 2017]. In Hole U1308A, planktonic foraminifera assemblages suggest that cold conditions occurred prior to and during H1.1. Toward the latter part of H1.1 moderate warming of surface water occurred and continued through the intervening period between H1.1 and H1.2 (Figure 8). We suggest that the warming during late H1.1 was related to the breakdown in stratification that occurred once the flux of icebergs was reduced. A slight cooling occurred during H1.2 followed by a return to warm conditions during the Bølling‐Allerød. The cold conditions during H1.1 are consistent with model results indicating widespread cooling of North Atlantic SSTs in the IRD belt during Heinrich events in response to freshwater forcing and expanded winter sea ice. Farther south at Site U1313 (41°N), Naafs et al. [2013] found evidence of rapid warming of 2–4°C associated with Heinrich events. At Site 976 in the Alboran Sea on the southern Iberian Margin (36°12.3′N, 4°18.7′W), HS1 is punctuated by a sudden brief warming (~4°C in less than eight centuries) at ~16 ka (Figure 12) [Martrat et al., 2014]. Measurement of bulk carbonate δ18O at Site 976 demonstrates that the warming preceded the decrease in bulk δ18O, which indicates the arrival of fine‐grained detrital carbonate during Heinrich Event 1 (Figure 12). Previous studies have shown that meltwater reached the western Mediterranean during Heinrich events [Sierro et al., 2005], which likely contained fine‐grained detrital carbonate from Hudson Strait [Hodell and Curtis, 2008]. Our results indicate a pattern of cooling and sea ice expansion in the high‐latitude North Atlantic during H1.1 when warming occurred in the subtropical eastern North Atlantic [Martrat et al., 2014]. These results support oceanic forcing and subsurface warming as a trigger of Heinrich Event 1 [Marcott et al., 2011; Naafs et al., 2013] with [Hulbe, 1997; Hulbe et al., 2004; Álvarez‐Solas et al., 2010, 2011] or without [Bassis et al., 2017] the existence of an ice shelf in the Labrador Sea. Figure 12 Open in figure viewer PowerPoint Martrat et al., 2014 18O of bulk carbonate at Site 976 in the Alboran Sea (this study), which we interpret as the delivery of fine‐grained detrital carbonate to the site during Heinrich Event 1. The yellow shading defines Heinrich Stadial 1 and gray shading HL 1.1. The bulk carbonate δ18O was analyzed twice as indicated by the filled circles (blue) and diamonds (red). Note the warming in SST preceded the decrease in bulk carbonate δ18O and peaks during H1.1. Sea surface temperatures (SSTs) recorded by alkenones [.,] relative to the δO of bulk carbonate at Site 976 in the Alboran Sea (this study), which we interpret as the delivery of fine‐grained detrital carbonate to the site during Heinrich Event 1. The yellow shading defines Heinrich Stadial 1 and gray shading HL 1.1. The bulk carbonate δO was analyzed twice as indicated by the filled circles (blue) and diamonds (red). Note the warming in SST preceded the decrease in bulk carbonate δO and peaks during H1.1.

4.5 Relationship of Heinrich Event 1 and Heinrich Stadial 1? It has been proposed that HE1 occurred when the polar North Atlantic was already cold and AMOC weakened [Bond et al., 1992; McManus et al., 1994]. Thus, HE1 was not the trigger for Heinrich Stadial 1 but increased freshwater forcing associated with HE1 may have weakened AMOC further. Figure 13 compares Ca/Sr from Site U1308 with Pa/Th from the Bermuda Rise [McManus et al., 1994] and Iberian Margin [Gherardi et al., 2005]. Overturning circulation in the western Atlantic Basin was suppressed throughout HS1 but was particularly weak during the latter half of HS1. The two detrital carbonate peaks at U1308 correspond with the minima in Pa/Th between ~16 and 15 ka. The Pa/Th record from the Iberian Margin suggests a somewhat different history in that the Pa/Th remains low during the first part of HS1 and increases later than the record from the Bermuda Rise [Gherardi et al., 2005, 2009]. On the Iberian Margin, the weakest period of overturning circulation occurs during the period when HL1.1 and HL1.2 were deposited (Figure 13). Figure 13 Open in figure viewer PowerPoint Toucanne et al., 2010 McManus et al., 2004 Gherardi et al., 2005 North Greenland Project Members, 2004]; (d) the oxygen isotope record of Hulu Cave, China [Wang et al., 2001 Wu et al., 2009 Southon et al., 2012 13C of atmospheric CO 2 from Taylor Glacier [Bauska et al., 2016 2 [Marcott et al., 2014 4 [Rhodes et al., 2015 Broecker and Putnam [ 2012 Comparison of (a) Ca/Sr of Site U1308 (filled black) and flux of Fleuve Manche sediment load (red,.,); (b) Pa/Th from the Bermuda Rise (green and blue) [.,] and Iberian Margin (red) [.,]; (c) the oxygen isotope record of Greenland [, 2004]; (d) the oxygen isotope record of Hulu Cave, China [.,.,, on time scale of.,]; (e) δC of atmospheric COfrom Taylor Glacier [.,]; (f) CO.,], and (g) CH.,] from the WAIS Divide ice core. The gray shaded areas denote the two‐stage subdivision of HS1 at 16.1 ka as suggested by]. Broecker and Putnam [2012] proposed a two‐stage subdivision of Heinrich Stadial 1 with a transition at ~16.1 ka. During the early stage of HS1 (18 to 16.1 ka), δ18O in Greenland ice cores was less negative indicating warmer temperatures than the latter stages of HS1 (Figure 13). The coldest conditions occurred during the latter half of HS1 (16.1 to 14.7 ka). The observation that H1 was not the cause of HS1 raises the question what was responsible for cooling and weakening of overturning circulation in early HS1 prior to the Heinrich event? Several studies show that IRD deposition prior to H1 in the central and eastern North Atlantic was derived mainly from European and Iceland ice sheets [Grousset et al., 2001; Peck et al., 2006]. Peck et al. [2006] suggested that instability and meltwater forcing by the Northwest European Ice Sheet temporarily weakened AMOC prior to HE1, illustrating that even modest ice sheets can have a disproportionate impact on deep‐water circulation if the ice sheet is close to the source of deep‐water formation. Toucanne et al. [2010, 2015] found a significant increase in sediment load to the Bay of Biscay at circa 20 ka that reflects the start of European deglaciation with a pronounced peak in Fleuve Manche discharge at 18 ka (Figure 13). The reduction in AMOC during the early part of HS1 may have been driven by earlier melt from the European ice sheet (~20–16.5 ka) [Peck et al., 2006; Toucanne et al., 2015], whereas the Laurentide played a more significant role during the latter half of HS1 from 16.5 to 14.7 ka.

4.6 Significance of Manganese Peaks The base of HL1.1 is sharp and underlain by dark sediment enriched in Mn (Figures 2, 3, and 5). Mn2+ is mobile under reducing conditions and precipitates as Mn (III, IV) under oxic conditions. It is typically enriched in oxic sediments with a concentration maximum near the surface where reduced Mn2+ migrates from reducing sediment below and is oxidized at the redox boundary that occurs near the depth of oxygen penetration into the sediments. Three Mn peaks occur in the upper 150 cm of Site U1308 indicating nonsteady state Mn redox cycling (Figure 5) [Burdige, 1993]. We suggest that the two older peaks represent relict Mn‐rich layers that formed at former redox boundaries, whereas the uppermost peak is actively forming today. Dissolution of the older Mn peaks may have been retarded by the deposition of HL1, which impeded upward diffusion of Mn2+ because of its high density and low porosity. Manganese peaks are often found at glacial‐interglacial transitions suggesting a change in redox conditions [Mangini et al., 1990, 1991; Jaccard et al., 2016; Gottschalk et al., 2016]. Because Mn2+ migrates in the sediments, the age of the sediment hosting the Mn peak provides a maximum age of the redox change [Mangini et al., 2001]. We suggest that the Mn peak associated with HL1.1 formed by a shoaling of the redox boundary in response to lowered oxygen and diminished ventilation of deep water. Similarly, the oldest Mn peak formed sometime during or after the Last Glacial Maximum. Both peaks are consistent with very old deep‐water reservoir ages (between ∼2250 and ∼3400 years) in the North Atlantic during the Last Glacial Maximum and HS1 [Skinner et al., 2014] and with lower oxygen concentrations of deep water on the Iberian Margin especially during HS1 [Hoogakker et al., 2015]. In addition to bottom‐water oxygenation, several factors can also influence Mn accumulation in marine sediments including changes in organic matter flux and sediment accumulation rate. Certainly, the deposition of the Heinrich layers resulted in large changes in sedimentation rate and we cannot rule out that the enrichment in Mn is a diagenetic consequence of changing sediment or organic fluxes [Thomson et al., 1996, 1999] rather than changes in deep‐water oxygenation.