Stable isotope evidence for glacial lake drainage through the St. Lawrence Estuary, eastern Canada, ∼13.1–12.9 ka

Quaternary International 260 (2012) 55e65

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Stable isotope evidence for glacial lake drainage through the St. Lawrence Estuary, eastern Canada, w13.1e12.9 ka T.M. Cronina, *, J.A. Rayburnb, J.-P. Guilbaultc, R. Thunelld, D.A. Franzie a

926A US Geological Survey, Reston, Virginia 20192, USA Department of Geological Sciences, State University of New York, New Paltz, NY 12561, USA c Musée de Paléontologie et de l’Évolution, Montréal, Québec H3K 2J1, Canada d School of the Earth, Ocean and Environment of South Carolina, Columbia, South Carolina 29208, USA e Center for Earth and Environmental Science, State University of New York, Plattsburgh, NY 12901, USA b

a r t i c l e i n f o

a b s t r a c t

Article history: Available online 10 September 2011

Postglacial varved and rhythmically-laminated clays deposited during the transition from glacial Lake Vermont (LV) to the Champlain Sea (CS) record hydrological changes in the Champlain-St. Lawrence Valley (CSLV) at the onset of the Younger Dryas w13.1e12.9 ka linked to glacial lake drainage events. Oxygen isotope (d18O) records of three species of benthic foraminifera (Cassidulina reniforme, Haynesina orbiculare, Islandiella helenae) from six sediment cores and the freshwater ostracode Candona from one core were studied. Results show six large isotope excursions (w0.5 to >2&) in C. reniforme d18O values, five excursions in H. orbiculare (<0.5 to w1.8&), and five smaller changes in I. helenae (<0.5&). d18O values in Candona show a 1.5e2& increase in the same interval. These isotopic excursions in cooccurring marine and freshwater species in varve-like sediments indicate complex hydrological changes in the earliest Champlain Sea, including brief (sub-annual) periods of complete freshening. One hypothesis to explain these results is that multiple abrupt freshwater influx events caused surface-tobottom freshening of the Champlain Sea over days to weeks. The most likely source of freshwater would have been drainage of the Morehead Phase of glacial Lake Agassiz, perhaps in a series of floods, ultimately draining out the St. Lawrence Estuary. Ó 2011 Elsevier Ltd and INQUA. All rights reserved.

1. Introduction During the last deglaciation the substantial freshwater flux from the ice sheets to the oceans was occasionally punctuated by abrupt (annual to decadal) proglacial lake drainage events, called catastrophic floods (Teller et al., 2002), triggered by threshold changes through glacio-isostatic rebound, outlet incision and/or ice-sheet margin retreat (Meissner and Clark, 2006). Geological evidence and climate modeling studies suggest these freshwater discharges can trigger changes in Atlantic Meridional Overturning Circulation (AMOC) and climate reversals such as the Younger Dryas (YD, Broecker et al., 1989; Manabe and Stouffer, 1995, 1997; Clark et al., 2001), the Preboreal Oscillation (PBO, Björck et al., 1996; Fisher et al., 2002; Cronin et al., 2008), and the 8.2 ka event (Barber et al., 1999; LeGrande et al., 2006). Broecker et al. (1989) proposed that rerouting of Lake Agassiz, the largest North American proglacial lake (13e8 ka, Teller et al., 2002), from a southward, Mississippi River/Gulf of Mexico route, to an eastward St. Lawrence * Corresponding author. Fax: þ703 648 6953. E-mail address: [email protected] (T.M. Cronin). 1040-6182/$ e see front matter Ó 2011 Elsevier Ltd and INQUA. All rights reserved. doi:10.1016/j.quaint.2011.08.041

Estuary/North Atlantic Ocean route led to YD cooling. Several studies support this hypothesis with both geomorphological (Teller, 1990; Teller et al., 2002; Teller and Leverington, 2004) and paleoceanographic evidence (Marchitto and Wei, 1995; Carlson et al., 2007; Rayburn et al., 2011). However, a number of studies have also questioned this hypothesis (Rodrigues and Vilks, 1994; de Vernal et al., 1996; Moore et al., 2000; Lowell et al., 2005, 2009; Teller et al., 2005; Broecker, 2006; Fisher and Lowell, 2006; Fisher et al., 2008). An alternative northwestern drainage route through the Mackenzie River into the western Arctic Ocean is possible based on field evidence (Lowell et al., 2005; Murton et al., 2010), paleoceanography (Poore et al., 1999), and modeling (Tarasov and Peltier, 2005; Smith and Gregory, 2009). Rayburn (1997), however, found it unlikely that the lake reached the northwestern outlet before or during the YD based on mapped northwestern strandlines of Lake Agassiz (see also Fisher and Lowell, 2012; Schell et al., 2008). Moreover, recent modeling studies suggest that a freshwater pulse from a lake outburst might be confined to the coastal boundary current system with minimal effect on subpolar areas of deep ocean convection (Condron and Winsor, 2011).

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Much of the controversy about glacial lake drainage stems from uncertainty in the ages, volumes, durations and pathways for lake drainage events related to uncertainty in Laurentide Ice Sheet (LIS) margin positions (Fisher and Lowell, 2006; Fisher et al., 2008; Lowell and Fisher, 2009) and the low temporal resolution and complex isotopic signature of paleoceanographic proxies (Flower et al., 2004; Hillaire-Marcel et al., 2008). This paper presents oxygen isotope (d18O) records from freshwater ostracodes and benthic foraminifera from proglacial Lake Vermont and the Champlain Sea sediments in the Champlain and St. Lawrence Valleys (CSLV) deposited between w13.2 and 10 ka, with emphasis on events around 13 ka. The approach assumes that, if sudden (years to decades) discharges of large volumes of glacial lake water drained eastward from lakes such as Agassiz into the CSLV, they would alter the d18O composition of Lake Vermont and Champlain Sea water, which would be reflected in the shells of calcareous microfaunas. If Lake Agassiz drained through the Great Lakes region into Lake Vermont, the d18O values of the freshwater ostracode Candona (d18OCand) should reflect the mixing of the two lake waters. Agassiz discharge into the Champlain Sea should influence the d18O values of benthic foraminifera (d18Oforam) by mixing enriched (0 to þ2&) marine and depleted (12 to 20&) lake water (Hillaire-Marcel, 1988; Cronin et al., 2008). As shown below, the results suggest an eastward drainage route for glacial Lake Agassiz water near the onset of the Younger Dryas. Testing the eastward drainage hypothesis in the CSLV region has several unique advantages: 1) the salinity effects of freshwater influx into the semi-isolated basin can be simulated using an estuarine circulation model and compared to proxy records (Katz et al., in press); 2) proglacial sediments were deposited at high sedimentation rates (w0.1e2 cm y1) (Rayburn et al., 2011) possibly preserving evidence for abrupt hydrological events; 3) the region is closer to North American glacial lakes than the more distal regions typically used to evaluate freshwater discharges, such as the St. Lawrence Estuary (de Vernal et al., 1996; Carlson et al., 2007), the Gulf of Mexico (Marchitto and Wei, 1995; Flower et al., 2004), and the Labrador Sea (Hillaire-Marcel et al., 2007, 2008); 4) the stratigraphy and chronology of the Laurentide Ice Sheet (LIS) margin in the region is well known (Hillaire-Marcel and Occhietti,

1977; Parent and Occhietti, 1999; Occhietti et al., 2001a; Occhietti, 2007; Dyke et al., 2002, 2003; Rayburn et al., 2005, 2007, 2011; Franzi et al., 2007); and 5) prior studies have shown the value of isotopic and faunal proxies for regional paleo-hydrological reconstruction (Cronin, 1977; Hillaire-Marcel and Occhietti, 1977; Corliss et al., 1982; Hillaire-Marcel, 1988; Rodrigues and Vilks, 1994; Cronin et al., 2008). With the exception of Corliss et al. (1982), this is the first study of Champlain Sea foraminiferal stable isotopes and, to the authors’ knowledge, the first study of stable isotopic analyses of co-occurring freshwater ostracodes and marine foraminifera from post-glacial sediments. 2. Regional setting The CSLV region lies north and east of the Adirondack Mountains of New York, north and west of the Green Mountains of Vermont, and south of the Laurentian Highlands in Canada. The region inundated by proglacial lakes and the Champlain Sea, the location of glacial Lakes Agassiz and Algonquin, the LIS margin w11.5e11.0 ka and possible lake drainage routes are shown in Fig. 1 along with the study sites described below (Table 1). Table 2 lists the proglacial lake drainage events in the region including the w13 ka events focused on below. Fig. 2 shows the informal phases of the proglacial water bodies of the northern Champlain Valley (Fig. 2a) from Franzi et al. (2007) and Rayburn et al. (2011). In the Champlain Valley, the typical lithological sequence is as follows: 1) subglacial deposits (till), overlain by rhythmically-laminated sediments (varves, laminated clays) deposited in glacial Lake Vermont (Coveville and Fort Ann Phases) (Fort Ann Phase is equivalent to Lake Candona (Parent and Occhietti, 1988; Parent and Occhietti, 1999; Ross et al., 2006) and Lake St. Lawrence (Rodrigues, 1992) in Canada); 2) glaciolacustrine and glaciomarine laminated muds and silts (Cummings and Occhietti, 2001) deposited during the Lake Vermont to the Champlain Sea (LV-CS) transition (Marine Phase I, Freshwater, and Transitional Phases); 3) silts and sands deposited during the Marine Phase II of the Champlain Sea (Marine Maximum and Regressive Phases).

Fig. 1. Map showing location of the Champlain Sea in the Champlain and St. Lawrence Valleys and glacial Lakes Agassiz and Algonquin (from Teller et al., 2002; Teller and Leverington, 2004). Major outlets for freshwater discharge to the ocean (St. Lawrence, Hudson, Mississippi, Mackenzie) and from Agassiz and Algonquin to the Champlain Sea (North Bay, Port Huron) are also shown. Dashed line is approximate Laurentide Ice Sheet margin position w11.5 ka from Dyke et al. (2002, 2003).

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Table 1 Sediment cores and outcrops used to study Champlain-St. Lawrence Valley Younger Dryas record. Region Champlain Champlain Champlain Champlain Champlain Champlain Canada Canada Canada Canada Canada Canada Canada Canada

Valley Valley Valley Valley Valley Valley

Site

Source

Latitude (N)

Longitude (W)

Beekmantown 1 Beekmantown 2 Peru 3/4 Marina 1 Marina 2 Salmon River Ile Perrot Ste-Philomène (Mercier) St-Alban Verchères Ste-Monique-de-Nicolet St-Roch-de-Richelieu St-Césaire St-Joseph-du-Lac

Rayburn et al., 2011 Rayburn et al., 2011 Rayburn et al., 2011 Rayburn et al., 2011 Rayburn et al., 2011 Rayburn et al., 2011 Guilbault, 1993 Guilbault, 1989 Guilbault, 1993 Guilbault, 1993 Guilbault, 1993 Guilbault, 1989 Guilbault, 1989 Guilbault, 1980

44.76 44.76 44.58 44.68 44.68 44.64 45.4 45.27 46.72 45.78 46.15 45.93 45.41 45.55

73.47 73.47 73.45 73.44 73.44 73.47 73.97 73.74 72.08 73.35 72.52 73.17 73.02 74.04

Fig. 2 also gives the conventional stratigraphy of the Champlain Sea based on macrofaunas (Fig. 2b), microfaunal zones from various authors (Fig. 2c and d), and phases of glacial Lakes Agassiz and Algonquin (Fig. 2e, f). The ages of faunal zones and age relationships between Champlain Valley and western lakes post-glacial deposits are approximate due to age uncertainty. The primary focus here is on the LV-CS transition (Marine Phase I, Freshwater, Transitional Phases), which occurred when the LIS, which dammed Lake Vermont, retreated north past the Warwick, Quebec area (Occhietti et al., 2001b). Lake level dropped w20e60 m depending on region (Parent and Occhietti, 1988; Rodrigues and Vilks, 1994; Rayburn et al., 2005), draining through the St. Lawrence Estuary into the North Atlantic, as marine water entered the isostatically depressed basin almost simultaneously. This transition, dated at 13.1 ka  0.2 ka, involved complex hydrological changes at the onset of the Younger Dryas climate reversal, which are discussed below.

3. Material and methods 3.1. Sediment cores Cores and outcrop sections used in this study are listed in Table 1 and shown in Fig. 3. The eight sites in eastern Canada, studied for foraminifera by Guilbault (1989, 1993), were used to construct a low-resolution isotopic record of the entire w3500-y long marine Champlain Sea episode. Six cores from four sites in the Champlain Valley of New York were cored in August 2004 and June 2006 using a mobile drill rig. Sites were chosen based on their proximity to natural exposures of the LV-CS transition (Franzi et al., 2007) and their stratigraphy, chronology and

Comments

Two boreholes F-1 and F-3 Two boreholes F1 and F2 800 m apart

micropaleontology are discussed in depth by Rayburn et al. (2011). 3.2. Micropaleontological and isotope analyses Calcareous benthic foraminifera occur commonly in CS sediments (Cronin, 1979; Corliss et al., 1982; Guilbault, 1989; Rodrigues, 1992; Cronin et al., 2008; Rayburn et al., 2011). Specimens of Haynesina orbiculare and Elphidium excavatum forma clavata (henceforth E. excavatum) from the eastern Canada sites at sample spacing ranging from 10 to 20 cm at St-Césaire and 1e3 m at other sites were analyzed for stable isotope composition. For the Champlain Valley cores, five to twenty shells of four species (Cassidulina reniforme, Islandiella helenae, H. orbiculare, and E. excavatum) from the >63 mm size fraction were analyzed at 1e2-cm intervals in the LV-CS transition interval and 2e10-cm intervals from the rest of the sequence. Stable isotope analyses were also conducted on adults and instars of the ostracode Candona subtriangulata from LV-CS transition sediments in the Beekmantown-1 and 2 cores from the northern Champlain Valley. Species identifications for the Canadian sites are from Guilbault (1989, 1993). All sediment processing and species identification for the Champlain Valley cores was conducted at USGS labs in Reston, Virginia and isotopic analyses were performed at the University of South Carolina, Columbia, South Carolina, stable isotope laboratory on a GV IsoPrime stable isotope ratio mass spectrometer (IRMS) equipped with a Multicarb sample preparation system. All isotope data are presented in per mil (&) d-notation relative to the Vienna Peedee belemnite (VPDB, NBS-19). Replicate analyses of this standard yielded an analytical precision of 0.08&. Data are available at the World Data Center for Paleoclimatology, http://www.ncdc.noaa. gov/paleo/paleo.html.

Table 2 Champlain-Hudson-St. Lawrence Valley glacial lake abrupt discharge events. Event

Age ka cal

Glacial Lake

Abandoned lake phase

Volume km3

Route

Estimated salinity excursiona

Reference

1

w13.3

Hudson

NA

Rayburn et al., 2005, 2007

w13.2 13.1e12.9 12.9

Main Iroquois & Coveville Frontenac Fort Ann Lockhart

700

2 3 4

Iroquois & Vermont Iroquois Vermont Agassiz

2500 1500 9500

Hudson St. Lawrence St. Lawrence

5 6

11.2e11.5 10.6-<10

Algonquin Agassiz (several events)

Main Algonquin Emerson-Nipigon

1000e2500 2500->5000

St. Lawrence St. Lawrence

NA 0 to >25 0 to 30; fluctuating 15 to 25 5 to 15

Rayburn et al., 2005, 2007 Rayburn et al., 2005, 2007, 2011 Teller and Leverington, 2004, Rayburn et al., 2011; this paper Cronin et al., 2008 Teller and Leverington, 2004

a

Baseline flow w56000 cubic m/sec integrates meltwater and flow through rivers into basin, includes rainfall, river discharge (Rayburn et al., 2005).

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T.M. Cronin et al. / Quaternary International 260 (2012) 55e65

a

b

c

d

e

f

Fig. 2. Postglacial stratigraphy and faunal zones of St. Lawrence-Champlain Valleys. 2a: Northern Champlain Valley stratigraphy (Franzi et al., 2007; Cronin et al., 2008). 2b. Champlain Sea phases (Elson, 1969; Cronin, 1977). 2c. Champlain Sea foraminiferal zones Pre-A through Post-C (Guilbault, 1989, 1993), zones F1B through F5 (Rodrigues and Vilks, 1994). Isotope record in Fig. 4 based on analyses of foraminifer from zones Pre-A through C. 2d. Southern Champlain Sea microfaunal zones (Cronin et al., 2008). 2e and 2f. Glacial Lake Agassiz (Teller et al., 2002; Teller and Leverington, 2004; Fisher et al., 2008) and Algonquin (Lewis and Anderson, 1989; Lewis et al., 1994; Moore et al., 2000) phases; correlations between Champlain Sea and glacial lake phases and faunal zones approximate.

3.3. Benthic foraminiferal and ostracodes stable isotopes There is a large literature on the ecology and isotope composition of the dominant Champlain Sea foraminiferal species, which today inhabit Arctic and subarctic shelves and fjord environments similar to glaciomarine conditions of the early Champlain Sea (Hald et al., 1994; Hald and Korsun, 1997; Korsun and Hald, 1998; Wollenburg and Mackensen, 1998; Sejrup et al., 2001, 2004; Polyak et al., 2002; Jennings et al., 2004; Scott et al., 2008). C. renifome is an infaunal species (Polyak et al., 2003; Ivanova et al., 2008) and secretes its shell in isotope equilibrium (Hald and Vorren, 1987) or at a slight (0.13&) offset from that of inorganic calcite precipitated in equilibrium (Austin and Kroon, 1996). H. orbiculare is a shallow infaunal or epifaunal species (Polyak et al., 2003) that shows a positive 1.5& offset in d18O values from equilibrium (Bauch et al., 2004). E. excavatum is an infaunal species that in some situations secretes its shell at positive 1.0& offset from equilibrium (Bauch et al., 2004), although elsewhere there is a negative offset (Polyak et al., 2003). Bauch et al. (2004) cautioned that E. excavatum d18O and d13C values are subject to vital effects in the Laptev Sea. d18O values for I. helenae (sometimes lumped with related species, Scott et al., 2008) are offset slightly (w 0.3e0.4&) from equilibrium (Knudsen et al., 2004; Slubowska-Woldengen et al., 2007). Despite possible vital effects, the isotopic composition of high-latitude benthic foraminiferal species has been used to reconstruct paleohydrography in high-latitude marginal seas,

including the North Sea (Bodén et al., 1997), Baffin Bay (Knudsen et al., 2008), the Barents Sea (Duplessy et al., 2001; Lubinski et al., 2001; Murdmaa et al., 2004), the Kara and Pechora Seas (Polyak et al., 2002, 2003), and off West Spitzbergen (Hald et al., 2004; Hald and Korsun, 2008). 3.4. Chronology Radiocarbon ages (14C) on Champlain Sea molluscs yield ages 350e1800 years older than ages from plant material from the same samples (Occhietti et al., 2001b; Occhietti and Richard, 2003; Richard and Occhietti, 2005; Cronin et al., 2008). Anomalous shell ages have been attributed to salinity variability (Rodrigues, 1988), old carbon reservoir effects (Hillaire-Marcel,1981,1988), vital effects for different species, or other unknown factors (Anderson, 1988; Rodrigues, 1992). Consequently, only radiocarbon ages on terrestrial plant material combined with varve counting, pollen stratigraphy, and ice sheet margin retreat rates are used for chronology in this study. Several lines of evidence date the LV-CS transition at the start of the Younger Dryas (discussed fully in Rayburn et al., 2011). The retreat of the ice margin from the northern Champlain Valley and the formation of Lake Vermont occurred prior to about 13.30 ka based on dates from musk ox bone and lake sediments. The age of the transition from lake to marine conditions between 13.12 and 12.85 ka (2-s age range) is based on a calibrated 14C date on wood from Champlain Sea deposits in the Marina 3 core, just above the LV-CS transition (1114 cm

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59

Fig. 3. Location of studied locations in eastern Quebec, core sites in the northern Champlain Valley, and in the Folger Trough in southern Lake Champlain (see Table 1 for study sites). Circles indicate sites used in this study.

core depth). The 4e5 m of sediment deposited during the transition, lying between basal till and the Marine Maximum Phase, represents w200 to 475 years of deposition at a mean rate of about 1e2 cm y1. The duration of the Freshwater Phase is estimated to be several decades to a century based on radiocarbon ages, varve and laminae thickness and number, and a regional ice margin retreat rate of 450 m y1. A wood date (12.39e11.98 ka) from 857 cm in the Salmon River core from the Marine Phase II suggests a slower mean sedimentation (w0.1e0.2 cm y1), consistent with those from CHIRP records and radiocarbon dates from the deep basin in modern day Lake Champlain (Freeman-Lynde et al., 1980; Cronin et al., 2008). The w13 ka age for the LV-CS transition in the study area is consistent with a calibrated date of 13.19 to 12.87 ka on plant

material from early CS sediments from near Mount St. Hilaire, Quebec (Richard and Occhietti, 2005), dates on plant material and CHIRP sonar records from southern Lake Champlain (Cronin et al., 2008), and numerous calibrated dates from throughout the basin (see review by Occhietti, 2007). 4. Results 4.1. Microfaunal assemblages The LV-CS transition is easily identified from microfaunal assemblages, which are shown for the Marina 3 core in Fig. 4. Rare but well-preserved benthic foraminifera, usually C. reniforme, first

60

T.M. Cronin et al. / Quaternary International 260 (2012) 55e65

Phase

Depth Foram & Candona Density

Polymorphinids

C. reniforme

I. helenae

H. orbiculare

Marine Phase II

1000

Regressive 1020

a

b

c

d

e

1040

Marine Max

Forams

1060 1080 1100

Transitional

1120

Freshwater Phase

1140 1160 1180

Marine Phase I

Candona

1200 1220 1240 0

1

2

3

4

0

0.5

1

1.5

2

0

3

6

9

12 15

0

0.5

1

1.5

2

0

0.5

1

1.5

2

Specimens/gram Fig. 4. Calcareous microfossil stratigraphy from Beekmantown 2 core showing typical faunal sequence during Lake Vermont-Champlain Sea transition (modified from Cronin et al., 2008). 3a. Density of freshwater ostracode Candona and marine benthic foraminifera, 4b-e. relative frequencies of Polymorphinids, Cassidulina reniforme, Islandiella helenae, Haynesina orbiculare. Earliest marine sedimentation (Marine Phase I (1250e1153 cm core depth) is earliest evidence for marine conditions (rare C. reniforme and I. helenae, salinity ¼ w25e30), sometimes mixed with the ostracode Candona subtriangulata (freshwater). Freshwater Phase (1152e1141 cm) brief return to freshwater environments (dominant Candona). Transitional Phase (1140e1099 cm) (polymorphinid zone). Marine Phase II marine maximum (1095e1040 cm), marked by the first occurrence of Stainforthia, dominant C. reniforme and I. helenae. Regressive Phase (<1040 cm) marked by dominant H. orbiculare. The microfaunal sequence, with minor variations, is observed in all core sites.

appear in laminated muds and sands of Marine Phase I (Fig. 4a). These sediments represent the earliest influx of marine water during the post-glacial interval. The low density of foraminiferal shells is attributed to high sedimentation rates, and also perhaps to low productivity due to the absence of light penetration, absence of annual marine diatom blooms and shortened period of adequate salinity. Sediments containing C. subtriangulata but no foraminifera constitute the Freshwater Phase, which follows the Marine Phase I. The Freshwater Phase is followed by the Transitional Phase, a zone containing both foraminifera and C. subtriangulata. A precise boundary between the Freshwater and Transitional Phases is difficult to define. For example, in the Beekmantown-1 core there are some foraminifera found with Candona and a few H. orbiculare occur in this interval in the Salmon River core. These intermittent occurrences of C. subtriangulata and foraminifera are probably caused by high and variable sedimentation rates and alternating brief periods of freshwater and marine conditions (see discussion). The last stratigraphic appearance of C. subtriangulata coincides with a distinct horizon marked by the early establishment of polymorphinid foraminifera and the ostracode Cytheromorpha macchesneyi (Fig. 4b). These two taxa persist side-by-side over a certain stratigraphic interval but, very quickly and over an important part of their interval of co-occurrence, they are completely overwhelmed by large numbers of the Marine Phase II species (Islandiella, etc.). The interval dominated by polymorphinids and C. macchesneyi is thus usually very thin but nevertheless present in every core. Sediments of Marine Phase II contain first a diverse foraminiferal assemblage with abundant C. reniforme and I. helenae (Fig. 4c,d, Marine Maximum) and then a zone of dominant H. orbiculare (Fig. 4e). This faunal sequence is found at all sites in the northern Champlain Valley with minor differences due to sedimentation rates, core recovery, and local environmental conditions (e.g., proximity to river input, water depth). Moreover, assemblages containing both Candona and foraminifera have been reported from other regions of CS marine inundation (Cronin, 1977; Hunt and Rathburn, 1988; Guilbault, 1989; Rodrigues, 1992).

4.2. Oxygen isotopes 4.2.1. Composite Champlain Sea d18O curve A composite oxygen isotope curve for the entire Champlain Sea episode was constructed from sites in Eastern Canada as a context for the d18O record from the transition zone presented in the next section. Mean d18O values on E. excavatum from Guilbault’s (1989) faunal zones (Pre-A, A, A/B, B, C and Post-C, see Fig. 2c) from eight sites are shown in Fig. 5. These sites represent relatively deep parts of the Champlain Sea Basin with deposition at about 50e100 m water depth. Zone Pre-A (Candona-bearing) from Ile Perrot, south of Montreal, and Zone A (diverse foraminiferal assemblage, no Candona) correlate approximately with the Transitional and Marine Maximum Phases, respectively. Notable features of the curve include a >5& increase (from 4 to þ1.19&) from Pre-A to A faunal zones, a 6& decrease from zone A to zone B, and a total decrease of about 9& (from þ1.19 to 8.09&) from Zone A to Zone C (Marine Phase II). The lowest d18O value found was 12.01& at the top of Zone C at St.-Roch-de-Richelieu. In addition to the isotope record shown in Fig. 5, d18O values from E. excavatum and H. orbiculare were obtained from samples representing shallow water facies of the Champlain Sea, foraminiferal zones EA and EH of (Guilbault, 1989) (not shown). Although more limited than the deep-water record, these samples also show a longterm 2e5& decrease in both species as the Champlain Sea shoaled. At St-Césaire, for example, H. orbiculare d18O values show a total 7.3& decrease in a series of samples from zones A, to EH and then to EA. In general, an increase in d18O during the early Champlain Sea (from Pre-A to A) followed by a long-term decrease suggests an initial rise followed by a long-term decrease in salinity during the remainder of the Champlain Sea episode. This is consistent with salinity variability inferred from the foraminiferal assemblages (Guilbault, 1989). 4.2.2. Champlain Valley d18O curve (w13.1e12.0 ka) Fig. 6 shows the benthic foraminiferal and ostracode d18O curves from six cores from the northern Champlain Valley and the

T.M. Cronin et al. / Quaternary International 260 (2012) 55e65

and Salmon River cores. In general, it is quite striking that Marine Maximum d18O values for all three species converge at their highest values (þ2.5&). Furthermore, with the exception of the Salmon River core, the Regressive Phase shows a gradual isotopic decrease of 0.9e1.73& at Beekmantown-2, Marina-2, -3 and Peru-3 in C. reniforme, 3e4& in H. orbiculare, and 1.3e1.4& in I. helenae. The lowest d18O values observed during the Regressive Phase in northern Champlain Valley cores are smaller than those for eastern Canada cores (Fig. 5). This is probably because core locations in the Champlain Valley were isostatically uplifted above the sea level earlier than the Canadian locations, ending deposition roughly from 12.5 to 12.0 ka, while deposition at the Canadian sites continued until about 10.6 ka.

Foram Zones C

B

A/B

A

Pre-A

-16

61

Figs. 6, 7

-12

-8

-4

0

18

δ O(‰) 18

Fig. 5. Composite d O record from E. excavatum from eight Canadian sites (Table 1) representing the entire Champlain Sea episode. Pre-A zone (Ile-Perrot site) contains mixed foraminifera and Candona and depleted values (4&). Zone A, the marine maximum, had most enriched values (0e1&). Zones A/B through Zone C show progressive depletion of 8e13&. Boxes show interval covered in detail in Figs. 6 and 7 (lakeemarine transition in the Champlain Valley). Blue symbols are mean values for each sample; red symbols mean values for each foraminiferal zone. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

stratigraphic position of key markers used to correlate among the cores (red sediment layer, last Candona occurrence, polymorphinid zone). Although foraminifers are sparse in Marine Phase I, all three species have high d18O values (2e2.5&) in cores Marina-2, Marina3 and Beekmantown-2. Marine Phase I is followed by the Freshwater Phase in which foraminifera are absent (see Candona isotopes below). In the Transitional Phase, there are large variations in the d18O composition of C. reniforme and H. orbiculare, ranging from <0.5 to 2.6& depending on the site. The Beekmantown-2 site had the most complete record of the transition and shows several isotope excursions of 0.5 to >1.5& in C. reniforme d18O, fewer, smaller oscillations in H. orbiculare d18O (<0.5e1.5&), and small oscillations in I. helenae d18O values which remain near þ2.0 to þ2.5&. These differences may reflect habitat preference for different species, vital effects, or absence of specimens in key intervals. For example, I. helenae is the most intolerant species to salinity reduction among the major Champlain Sea species (Polyak et al., 2002, lumped with Islandiella norcrossi) and would not live in periods of very low salinity, however short. This may be why its d18O values are relatively invariant. The total increase in d18O values from the Transitional Phase to the Marine Maximum ranged from about 1& to >3.5& depending on the core site and the species. The increase was largest in the Peru

4.2.3. Candona d18O record The Beekmantown-2 core contained the most detailed record of the Freshwater and Transitional Phases, including the zone containing both Candona and foraminifera in laminated sediments from 503 to 454 cm core depth. Fig. 7 shows that d18OCand values at 508e503 cm core depth were about 13.3 to 13.8& followed by rapid increase to 11.5& at 469 cm, then stabilization near 12.2 to 12.8& at 467 to 454 cm. A 2& increase is also seen in Candona from the Beekmantown-1 core from 524 to 454 cm core depth. C. reniforme values in the Beekmantown-2 interval containing Candona (488-439 cm) show fluctuating values from þ0.1 to þ2.0& with six negative excursions reaching minimal values ranging from þ0.1 to 1.2&. A decrease of 1.19& in C. reniforme from 486 to 461 cm corresponds to a 1.99& increase in Candona at 478e469 cm. H. orbiculare also shows a wide range of values in this interval from a low of 0.73& (467 cm) to high of 2.5& (457 cm), although there are fewer excursions than for C. reniforme. I. helenae d18O values in the interval containing Candona ranged from 2.1 to 2.7& but no specimens were found in Beekmantown-2 from 450 to 467 cm. 5. Discussion Several lines of evidence suggest rapidly fluctuating salinities during the Lake Vermont-Champlain Sea Transition about 13 ka. Microfaunal assemblages. Mixed foraminiferal and Candona assemblages occur in the oldest Champlain Sea deposits in the northern Champlain Valley (Cronin, 1977) and many sites in eastern Canada (Guilbault, 1989). Hunt and Rathburn (1988) also found Candona-bearing sediments overlying foraminfieral beds in cores from modern Lake Champlain. As shown in Fig. 4 above and in Rayburn et al. (2011), mixed and alternating foraminifera and Candona-bearing assemblages confirm these patterns in the current study area. The ostracode Candona inhabits exclusively non-marine environments; CS benthic foraminifera inhabit marine to brackish conditions. Their co-occurrence seems to be ecologically impossible if indeed they were contemporaneous. Shell preservation and biostratigraphy throughout the Champlain Sea deposition suggest the Candona-foraminifera co-occurrence cannot be explained by local transport or reworking. Cummings and Russell (2007) reached the same conclusion in studies of the LV-CS transition interval in Ontario. The only reasonable explanation is that hydrological and salinity changes too rapid to be distinguished using microfaunas, at least in some cores, occurred near the time of the initial influx of marine water across all or most of the Champlain Sea basin resulting in the alternation and sometimes mixing of freshwater and marine species. Lithology. Rhythmically laminated muds deposited during the transition containing foraminifers and sometimes Candona contrast with underlying lacustrine varves and overlying massive marine muds. Ross et al. (2006) studied stratigraphic and geophysical

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Fig. 6. d18O records for three foraminiferal species (Cassidulina reniforme, Islandiella helenae, Haynesina orbiculare) from 6 cores in Northern Champlain Valley. Intervals highlighted for each stratigraphic phase of the Champlain Sea based on the faunal (Fig. 4) and isotopic records for each core including three key marker horizons (Candona Last Appearance, polymorphinid zone, red sediment layer). During Freshwater and Transitional Phases, the three species do not show similar isotopic patterns. During Marine Maximum all three species have their highest d18O values; during Regressive Phases there is gradual depletion in d18O for all three species. Radiocarbon age from Marina-3 dates the Marine Maximum at w13.1e12.8 ka. See Fig. 7 for detailed record of Freshwater and Transitional Phases.

a

b

-12 -13 -14 -15

340

340

Islandiella

Beek-2

290

Beek-1

H. Orbiculare C. reniforme

360

380

380

400

400

420

420

340

390

Marine Maximum 440

440

460

460

480

480

500

500

Core Depth (cm)

Core Depth (cm)

360

440

490

520 0.0

0.5

1.0

1.5

2.0

2.5

520 -11 -12 -13 -14 3.0

Fig. 7. Comparison of foraminiferal d18O record from Beekmantown 1 (Fig. 7a) with Candona d18O records from Beekmantown 1 and 2 cores (Fig. 7b, X-axis reversed). Note progressive 2& enrichment in Candona d18O in both cores with a brief 1& depletion event in Beek-1 and Beek-2 at w450 cm and 470 to 465 cm core depths, corresponding to depletion in H. orbiculare d18O values. The interval from 480 to 435 consists of silt/ clay couplets about 3 mm thick at the base which grade upwards to 1 mm at the top where there is a sand parting and a sharp transition to massive mud. Assuming these couplets to be annual is consistent with estimated sedimentation rates from other cores in the region (Rayburn et al., 2011).

records northwest of Montreal, especially in Chemin des Sources region, noting the difficulty distinguishing between glaciolacustrine and glaciomarine outwash fan sediments. They also found an erosional unconformity between Lake Candona and Champlain Sea sediments in the Pointe-aux-Sables section. When marine sedimentation begins at Pointe-aux-Sables, it is already in late CS time (10,459  40 on Hiatella arctica; Ross et al., 2006) and there are abundant Elphidium albiumbilicatum, as in Zone EA (Guilbault, unpublished data). Cummings and Occhietti (2001) described “energetic subaequeous fan deposits” near Quebec City consistent with deposition at high sedimentation rates. Cummings and Russell (2007) interpreted transitional sediments in the Vars-Wincester esker southeast of Ottawa as reflecting sediment-laden, glacial meltwater jeteplumes flowing into standing brackish water of the Champlain Sea. They hypothesized a rapidly fluctuating salinity front caused by astronomic (diurnal, seasonal), episodic (flood) forcing or dynamic meltwater changes along ice sheet margin. Stable isotopes. The isotope data provide additional insight into the nature of the lakeemarine transition. Although vital effects and habitat preferences might explain small interspecific differences in d18Oforam values, excursions of 1e2& in C. reniforme and H. orbiculare suggest large hydrological changes occurred over very short timescales. Global ice volume effects (Chappell and Shackleton, 1986; Mix, 1992) would have been negligible during the brief period of deposition during the LV-CS transition. Similarly, bottom water temperature remained near 0  C during the early Champlain Sea (de Vernal et al., 1996; Hillaire-Marcel, 1988; Hillaire-Marcel personal communication, 2007) and changes in d18Oforam of several per mil would require a 8e12  C warming of bottom waters, which is unlikely. Therefore, this supports HillaireMarcel’s (1988) conclusion that dilution by freshwater is the most likely explanation for d18O variability in CS carbonate fossils. Possible hydrological processes that might cause both the d18Oforam and microfaunal patterns include changes in precipitation-evaporation in the early CS basin, reduced or

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oscillating marine water inflow, or the influx of d18O-depleted freshwater from river discharge, Laurentide Ice Sheet meltwater, or glacial lakes. It is difficult to imagine that changes in precipitation, river discharge or evaporation would have been large enough to cause basin-wide isotopic and faunal changes on a decadal scale or less. Rapid decreases in marine water influx that might in theory be caused by a step-wise oscillatory nature to earliest marine inflow or to brief readvances or surges of the LIS margin surges near the Quebec City region cannot be totally discounted. However, stratigraphic and glacial geological evidence for dynamic LIS margin changes are lacking (Parent and Occhietti, 1999; Occhietti et al., 2001b; Occhietti, 2007). The most likely explanation involves either a rapid increase in melting rate from the LIS margin or freshwater influx from glacial lakes. Cummings and Russell (2007) postulated that LIS meltwater might have caused fresh-marine oscillations during the lakeemarine transition observed in the Vars-Winchester esker sequence in Ontario. However, the LIS margin was more than 150 km north of the northern Champlain Valley where similar hydrological changes occurred making this explanation unlikely, although the synchroneity of the changes in the two regions is not established. This leaves glacial lake water influx as the most likely cause for the rapid changes in oxygen isotopes. To evaluate this possibility, Champlain Sea d18Oforam values are considered in terms of the isotopic composition of potential marine and freshwater sources. The CS was stratified in terms of salinity and oxygen isotopic composition similar to the modern Belle Isle Strait region between Newfoundland and Labrador where fresher, isotopically depleted water overlie saline, isotopically enriched water. Hillaire-Marcel (1988) estimated that deep water (50e60 m) in the Champlain Sea had d18Owater of 0 to þ2& and salinities of 25e28 psu. It is estimated that the þ2& foraminiferal values discussed above correspond to salinities of w 30 psu for brief intervals during the Transitional Phase and during the Marine Maximum. There is a range of estimated d18Owater values from Lake Agassiz and Algonquin (Lake Huron Basin) derived from C. subtriangulata and other material from lake sediments in Lake Agassiz and the Great Lakes region (Last et al., 1994; Rea et al., 1994; Remenda et al., 1994; Dettman et al., 1995; Buhay and Betcher, 1998; Rodrigues and Lewis, 2000; Birks et al., 2007; Hillaire-Marcel et al., 2008; Breckenridge and Johnson, 2009). Last et al. (1994) obtained values of 15.3 to 16.4& VPDB (¼ 21 to 18 SMOW) from Candona d18O for the early Agassiz phase epilimnion and 24& for the hypolimnion. Birks et al. (2007) estimated the hypolimnion waters of the Lockhart (pre-YD) Phase of Lake Agassiz to be about 24& per mil VPDB based on measured cellulose and modeled porewater d18O and about 17 to 20& based on ostracode d18O. Cellulose d18O values from the Moorhead (YD age) Phase were about 15 to 19&. Birks also concluded that epilimnion values became increasingly enriched due to decreasing influence of LIS meltwater and increasing influence of precipitation and possibly evaporative effects. Hillaire-Marcel et al. (2008) measured d18OCand at 20.5& VPDB from lacustrine varved sediments underlying marine clays in southeast Hudson Bay. Assuming a lake bottom temperature of about 0  C, he estimated the isotopic composition of 25& VSMOW for late stage Agassiz. This value is only slightly more negative than estimated for Laurentide Ice Sheet (LIS) meltwaters during the last glacial maximum (w 25 to 35&, Flower et al., 2004; or 21e24&, Vetter et al., 2009). In summary, it is not possible to know the exact Lake Agassiz d18Owater values that might have entered the Champlain Sea at 13 ka, but the oscillating d18Oforam and d18OCand values during the LV-CS transition support the idea that freshwater influx altered the d18O of the host water body. Like the co-occurrence of foraminifera

63

and freshwater ostracodes in the transition zone, this evidence can be best explained by hydrological changes that occur more rapidly than the time interval represented by the 1e2 cm sampling interval. This would mean complete freshening of the basin occurred in less than a year, perhaps several times. Clearly no modern analog exists for a sub-annual scale shift from nearly fully marine conditions to freshwater and back again in a large semienclosed basin like the Champlain Sea. One can, however, envision a scenario in which a large enough volume of lake water entered the Champlain Sea rapidly, briefly freshening the basin from top to bottom by pushing eastward the saltwater wedge. Teller et al. (2002) estimated that Agassiz floods discharged from 1600 km3 (Pas Phase w9.2 ka) to 9500 km3 (Herman Phase w13ka) of freshwater over a year or less. Modeling studies carried out in conjunction with the current study show that Champlain Sea salinity responds (decreases) to the introduction of these volumes of freshwater within days and responds to the cessation of flooding events (salinity increase) within just weeks (Katz et al., in press). Thus, it is not unreasonable to expect faunal and isotopic records from sediments deposited during these extraordinary events to show such an enigmatic record of hydrological changes. Taken together, the faunal, isotopic and modeling results support the original hypothesis of Broecker, discussed in depth in Teller et al. (2002), that Lake Agassiz drained east through the St. Lawrence Estuary at the inception of the Younger Dryas, perhaps in a series of brief discharges. The impact on Champlain Sea salinity was swift (days) and immense (complete freshening), but near marine salinity returned quickly (weeks), such that it is unlikely there was a sustained effect on estuarine or surface ocean conditions in distal regions in the outer St. Lawrence Estuary or the Labrador Sea. Acknowledgments We appreciate the help of J. Dyszynski, M. Berke, M. MacNamara and R. Glazer with sample processing, J. Farmer with graphics, G. Cobbs with drilling, and C. Bernhardt, D. Willard and anonymous reviewers for helpful comments. Funded by USGS Global Change Program. Appendix. Supplementary material Supplementary data related to this article can be found online at doi:10.1016/j.quaint.2011.08.041. References Anderson, T.W., 1988. Late Quaternary pollen stratigraphy of the Ottawa Valley-Lake Ontario region and its application in dating the Champlain sea. In: Gadd, N.R. (Ed.), The Late Quaternary Development of the Champlain Sea Basin. Geological Association of Canada Special Paper, vol. 35, pp. 207e224. Austin, W.E.N., Kroon, D., 1996. Lateglacial sedimentology, foraminifera and stable isotope stratigraphy of the Hebridean Continental Shelf, northwest Scotland. In: Andrews, J., Austin, W., Bergsten, H., Jennings, A. (Eds.), The Lateglacial Palaeoceanography of the North Atlantic Margins. Geological Society of London, vol. 111, pp. 187e213. Barber, D., Dyke, A., Hillaire-Marcel, C., Jennings, A.E., Andrews, J.T., Kerwin, M.W., Bilodeau, G., McNeely, R., Southon, J., Morehead, M.D., Gagnon, J.-M., 1999. Forcing of the cold event of 8,200 years by catastrophic drainage of Laurentide lakes. Nature 400, 344e348. Bauch, H.A., Erlenkeuser, H., Bauch, D., Mueller-Lupp, T., Taldenkova, E., 2004. Stable oxygen and carbon isotopes in modern benthic foraminifera from the Laptev Sea shelf: implications for reconstructing proglacial and profluvial environments in the Arctic. Marine Micropaleontology 51, 285e300. Birks, S.J., Edwards, T.W.D., Remenda, V.H., 2007. Isotopic evolution of glacial Lake Agassiz: new insights from cellulose and porewater isotopic archives. Palaeogeography, Palaeoclimatology, Palaeoecology 246, 8e22. Björck, S., Kromer, B., Johnsen, S., Bennike, O., Hammarlund, D., Lemdahl, G., Possnert, G., Rasmussen, T.L., Wohlfarth, B., Hammer, C.U., Spurk, M., 1996.

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