FOP Field Trip Guide by:

Stephen F. Wright ¹ , Richard K. Dunn ² , George E. Springston ² , and Laurie D. Grigg ² 

 ¹ Department of Geography and Geoscience, University of Vermont, Burlington, Vermont

 ² Department of Earth and Environmental Sciences, Norwich University, Northfield, Vermont

Digital Story Map by: Peter Strand ³ 

 ³ Vermont Geological Survey, Montpelier, Vermont

We welcome you to the 85th annual reunion of the Friends of the Pleistocene, the first Friends’ field trip to northern Vermont since 2003. During most of the intervening 20 years our efforts have been directed towards detailed mapping of all 7.5-minute quadrangles within the Montpelier, Vermont 1:100,000 (30’ x 60’) sheet. This trip celebrates the compilation of the surficial geology across these 32 quadrangles. Feld Trip stops are selected to illustrate (1) glacial/hydrologic processes operating as the ice sheet thinned and retreated across this mountainous terrane, (2) glaciolacustrine processes occurring in the glacial lakes that formed during ice sheet retreat in the regions’ valleys, and (3) post-glacial processes and climate changes occurring across the area following the drainage of these lakes.

Use the following Story Map guide to revisit stops from the 2024 Field Trip. Click on an individual stop for additional information and photographs.

This field trip guide summarizes our current understanding of the glacial geology of north-central Vermont (Fig. 1). The surficial geology of the entire state was mapped during the 1950’s and 1960’s and that work compiled to make the “Surficial Geologic Map of Vermont” (Stewart and MacClintock, 1970). During that mapping effort geologists were tasked with mapping 15-minute quadrangles in a single field season which necessitated greater and lesser degrees of reconnaissance mapping. Their mapping efforts were also limited by the detail offered on the 15-minute topographic base maps available at the time. On the other hand, they had the advantage of working in a landscape that was more open and accessible and at a time when many active pits were supplying sand and gravel for interstate construction, many of which have been long closed. Also assembled at this time was an inventory of highway construction materials where the state surveyed every town to map the extent of sand and gravel deposits available for future highway construction projects (see  Vermont Geological Survey website  for downloads of these reports).

The impetus for this gathering of the Friends of the Pleistocene is the completion of detailed mapping of all 32 7.5- minute quadrangles within the Montpelier 1:100,000 sheet (Fig. 1). This mapping work began in the 1990's using paper quadrangle maps and pencils and evolved to utilize the latest phone-based, GPS-enabled mapping software on detailed LiDAR shaded-relief base maps. This mapping work has led to a better understanding of a wide variety of different glacial processes occurring as the Laurentide Ice Sheet flowed across the region at the Last Glacial Maximum and then subsequently thinned and retreated from the area. We outline below many of the ideas presented over the last 20 years in GSA talks, Vermont Geological Survey Reports that accompany individual quadrangle maps, and NEIGC field guides. The following sections of this field guide provide background to the field stops which are described as the last part of this guide.

Glacial till, erratics, and striations on the summits of the region’s highest mountains indicate that the ice sheet in northern New England was sufficiently thick to completely cover the mountains. The direction of ice flow is interpreted from glacial striations, rouches moutonnées, erratic dispersal fans, and most recently by crag and tail and streamlined till landforms visible in LiDAR imagery. While till fabrics were utilized extensively in the 1960’s to interpret several different regional ice flow directions occurring in separate tills (Stewart and MacClintock, 1969, 1970), this work has been largely discredited (see discussion in Wright, 2015b).

Using erratic dispersal fans emanating southeast and south from intrusive granite stocks in central Vermont, Larsen (1972) concluded that the Laurentide Ice Sheet had only flowed towards the southeast and south during the late Wisconsinan. Larsen (2003b) was also able to successfully trace a train of distinctive gneissic erratics from central Vermont in a NNW direction to their source, the St. Didace pluton in the Laurentian Mountains of Québec. A recent map showing the locations of distinctive orbicular granite erratics emanating from the one known outcrop in Craftsbury shows a very similar dispersal fan (Fig. 2).

A compilation of glacial striations observed in northern and central Vermont was published by Wright (2015b) and this compilation has been updated with new data to accompany publication of new Montpelier Surficial Geologic Map (Fig. 3). The most widely distributed set of glacial striations are aligned generally northwest-southeast across northern Vermont and are particularly well preserved along the high ridge lines of both the Green Mountains and lower elevation mountain ranges to the east (Northfield, Worcester, and Lowell Ranges, Fig. 3, 4). This striation set maintains a relatively uniform azimuth across north-central Vermont and is parallel with striations measured across a wide swath of south-central New Hampshire and southern Maine, areas immediately downglacier from north-central Vermont (Goldthwait in Antevs, 1922; Thompson and Borns, 1985) as well as in adjacent parts of southern Québec (Parent and Occhietti, 1999). The regional distribution of this striation set, particularly at high elevations, suggests a broad area of southeast ice flow existed at the LGM with ice flowing obliquely across the mountains in northern Vermont.

Figure 5 models the ice sheet profile across New England at a time when the ice sheet was flowing to the southeast at or near its maximum extent. The ice sheet profile is computed using the ground surface slope and a spreadsheet algorithm developed by Benn and Hulton (2010) set to maintain a uniform basal shear stress of 100 kPa, an average value. The calculated ice sheet surface is well over 2 km above the Green Mountain ridgeline and almost 3 km above the ground surface in the Champlain valley (right side of Fig. 5). Lower basal shear stresses would thin the modeled ice sheet profile whereas higher basal shear stresses would thicken the modeled profile. It should be noted that in contrast to lower elevation areas, exposure age dates from the highest elevation peaks in New England have consistently recorded old dates indicating relatively little bedrock was eroded from these peaks by the ice sheet, possibly because the ice sheet was cold-based in these areas (Bierman et al., 2015; Corbett et al., 2019; Halsted et al., 2022). The below model does not account for these “sticky spots” at the base of the ice sheet and, given their very small area across New England, they are unlikely to significantly affect the ice sheet profile.

In the many mountain valleys the youngest striations usually parallel those valleys (Fig. 3). The logical interpretation is that topography began influencing ice flow as the ice sheet thinned, a common observation along the margin of modern ice sheets. In most areas ice flow shifted from southeast to south as most of the valleys in the region are aligned parallel to the north-south trend of the mountains (Fig. 3). This is particularly evident in the Champlain valley west of the mountains and in the broad area east of the mountains centered on the Connecticut River valley. Large scale exceptions to this are the Lamoille and Winooski river valleys which both cut completely across the mountains and effectively funneled ice from the Champlain Valley ESE across the mountains. On the west side of the mountains glacial striations converge into these valleys and the aerial extent of this convergence is an indication of the area of ice in the Champlain Valley that was contributing to these outlet glaciers (Fig. 3). A fuller description of the glacial lake history follows in a later section, but these outlet glaciers were responsible for damming those major valleys cutting across the mountains producing large glacial lakes east of the mountains (Larsen, 1972, 1987).

 From Appalachian Gap south to Killington, a younger set of striations occurs along the Green Mountain ridge line and is oriented from northeast to southwest (Figs. 3, 4). These striations indicate that an abrupt shift in ice flow occurred from southeast to southwest resulting in ice now flowing across the mountains into the Champlain valley (Ackerly and Larsen, 1987; Wright, 2015b). Wright (2015b, 2017) attributed this change in ice flow to the initiation of an ice stream in the southern Champlain and northern Hudson River valleys. Abundant north-south striations, crag and tail structures, and streamlined bedforms with length to width ratios exceeding 10:1 within the Champlain valley all indicate that fast north to south ice flow occurred parallel to this large-scale topographic trough (Wright, 2015b).

The physiography of much of central-northern Vermont is dominated by the watersheds of the westward flowing Winooski and Lamoille rivers as they cut perpendicularly across the NNE-SSW strike of folded, early to middle Paleozoic metasedimentary rocks (Fig. 1). Major tributaries to the Winooski and Lamoille rivers flow parallel to bedrock strike. In this section we outline the evolution of glacial lakes in these valleys concentrating on areas east of the Green Mountains. We note, however, that our work in this region has not fundamentally altered the glacial lake history first sketched by Merwin (1908) and later deciphered in some detail by Larsen (1972, 1987). Larsen (1972, 1987) also presents excellent reviews of early work in the Winooski drainage basin leading to our current understanding of this history.

During glacial retreat across central Vermont the ice margin dammed the north-draining tributary valleys of the Winooski; from west to east the Mad River, the Dog River, and the Steven’s Branch. These valleys filled with meltwater and the subsequent lakes drained over southern thresholds (Fig. 6; Larsen, 1972, 1987). Deltas located in each of these valleys are graded to the different lake levels, confirming the relative elevations of these lakes (Dunn et al., 2007; Larsen, 1972, 1987). Glacial lake projections are largely from Springston et al. (2020) and use an isostatic tilt of 1.15 m/km to N21W.

These early lakes grew in size as the ice continued retreating northwards until the ice margin approached the low- lying Winooski River valley allowing the small lakes to coalesce and form a single larger lake, Glacial Lake Winooski, with a spillway across the lowest 276 m drainage divide (Fig. 7; Larsen, 1972, 1987). Recent mapping of this spillway area has shown that the lake outlet elevation originally identified by Larsen (279 m (915 ft); 1972, 1987) is covered by two Holocene alluvial fans and the outlet elevation is actually a few meters lower in elevation (276 m) and several hundred meters farther south (Wright, 2022b). The lake continued expanding north following the retreating ice front and eventually inundated both the Winooski and Lamoille River drainage basins (Fig. 8).

The maximum northwestern extent of the lake is known from the limit of Glacial Lake Winooski sediments mapped in the Lamoille River valley (Wright, 2001). Drainage of the lake remained at the Williamstown Gulf spillway in the southeast. During this time the water discharging over that spillway was sourced not only from meteoric water across both drainage basins, but glacial meltwater from rapidly thinning and retreating ice east of the Green Mountains and meltwater routed from the Champlain valley across the mountains via the Lamoille and Winooski River valleys. No discharge estimates have been made, but considerable volumes of water must have been flowing across this spillway.

Continued ice retreat down the Winooski valley eventually opened two new closely spaced outlets in the vicinity of the Huntington River (Fig. 9). The first of these, the Gillette Pond outlet, lowered Glacial Lake Winooski by ~90 m rapidly draining ~375 km3 of water from the lake basin (Wright, 2018a). It’s unclear whether the ice dam failure was via (1) a subaerial breach (water flowing over the ice sheet), (2) a subglacial tunnel, (3) flotation of the ice margin, or some combination of these processes. Regardless, the ice dam failure was likely catastrophic as friction from water leaking out of Glacial Lake Winooski rapidly melted the ice sheet surrounding the outlet passageway. The “new,” lower-elevation lake is referred to as Glacial Lake Mansfield 1 which drained, via Glacial Lake Huntington in the Huntington River valley into Glacial Lake Vermont (Fig. 9; Larsen, 1972, 1987). A well-preserved delta and channel occurs where the Gillette Pond outlet enters the Huntington River valley (Fig. 10).

Only several hundred meters of additional ice retreat allowed Glacial Lakes Mansfield and Huntington to merge and utilize a common outlet at the Hollow Brook threshold ~26 m lower than the Gillette Pond outlet (Fig. 9). This lower elevation stage is referred to as Glacial Lake Mansfield 2 (Larsen, 1972, 1987). When the ice sheet retreated into the Champlain valley, Glacial Lake Mansfield 2 drained through an outlet in Williston and the Winooski River valley became an arm of Glacial Lake Vermont (Fig. 11). Glacial Lake Vermont underwent several significant drops in elevation during its history (Chapman, 1937; Rayburn et al., 2011; Rayburn et al., 2005; Van Hoesen et al., 2016), but the shoreline changes associated with these different lake stages in the Winooski River valley are limited due to the steep topographic gradients (Fig. 11).

Typical of areas inundated by glacial lakes, large parts of the Winooski River basin below the elevation of Glacial Lake Winooski are blanketed with a variety of glaciolacustrine sediments (Wright et al., 2023a). In most areas these sediments occur in a fining-up sequence that records the retreat of the ice front, the source of most of the sediments, progressively farther from the locus of sediment deposition (Larsen, 1987). The sections typically begin with coarse esker sediments which transition into subaqueous fan deposits (aka ice-proximal deposits). These then transition to thick sequences of very fine and fine sand and then to varved silt and clay, i.e. ice-distal deposits. Our mapping indicates that in the shallow arms of Glacial Lake Winooski rapid sedimentation adjacent to the retreating ice front likely completely filled those parts of the lake with sediments. The significant ~90 m drop in lake level from Glacial Lake Winooski to Glacial Lake Mansfield 1 triggered rapid erosion of earlier-deposited sediments and their later redeposition as deltas along the margins of the lower-elevation lake.

During deglaciation, many small, unnamed proglacial lakes formed in the irregular topography of central Vermont. Similarly, during a short-lived and probably localized readvance of the ice margin, small proglacial lakes were impounded in several valleys. On Sunday we will visit depositional sequences in two of these lakes, glacial Lake Nasmith and glacial Lake East Barre (Dunn and Springston, 2019), which formed in north-flowing tributaries of the upper Winooski River.

The timing of ice advance across New England has been difficult to asses. Based on the youngest OSL ages of cave sediments in the Champlain valley Munroe and others (2016) concluded that the Laurentide Ice Sheet did not advance south of the St Lawrence lowlands into the Champlain valley until early in MIS-2, ~35 ka BP. The ice sheet subsequently advances across New England reaching the terminal moraine by ~25 ka BP (Balco et al., 2002; Corbett et al., 2017) a time span of ~10,000 years.

Our understanding of the timing of ice sheet retreat across northern Vermont comes from both the glacial lake varve record (the North American Varve Chronology; Ridge et al., 2012) and from exposure age dating. We begin with the general observation that the ice sheet east of the Green Mountains retreated northward much earlier than the ice sheet in the Champlain valley. We know this because Glacial Lake Winooski expanded well north of the Lamoille River valley at the same time that ice much farther south in the Champlain valley was sufficiently thick to occupy and dam the Winooski River forming the lake (Fig. 8).

Building on early work by Antevs (1922, 1928), Ridge and others (2012) successfully correlated and dated numerous sections of varves along the Connecticut River valley establishing the North American Varve Chronology (NAVC) and dating the retreat of the ice sheet up the valley. Figure 12 transcribes Ridge’s (2016) dated ice margin positions along a portion of the Vermont/New Hampshire border. During this period of time the ice sheet was retreating at a rate of ~300 m/yr, markedly faster than its 80-90 m/yr retreat rate across Massachusetts’ border with Vermont and New Hampshire (Ridge et al., 2012).

The extension of the 14.1 and 13.9 kyr ice margin positions across the Winooski drainage basin is based on the correlation of Glacial Lake Winooski varve sections originally measured by Antevs (1928; Fig. 13). More recently, three new sections of varved Glacial Lake Winooski sediments have been measured and a fourth, originally mapped by Antevs, has been remeasured and extended (Fig. 13). Specifically these sections are the Muzzy Brook (Larsen et al., 2001), Montpelier (Springston and colleagues, unpublished), Wrightsville Reservoir (Wright and students, unpublished), and Waterbury Reservoir Sections (Wright, in Larsen et al., 2003b). Figure 14 shows how these sections are correlated with Antevs’ (1928) Winooski Valley varve compilation and the North American Varve Chronology.

The Waterbury Reservoir section (Fig. 14, red line) clearly records when Glacial Lake Winooski partially drained. An abrupt change from silt/clay varves to thick fine sand/silt layers occurs at North American Varve 6976 which was deposited ~13,844 yr BP based on dating of the North American Varve Chronology (Thompson et al., 2017). Calibration of the NAVC has an uncertainty of at least 100 years (Thompson et al., 2017).

The increase in both grain size and layer thickness occurred when the ice dam in the Winooski Valley retreated far enough down-valley (WNW) for water in Glacial Lake Winooski to rapidly escape to the Champlain valley via the Gillette Pond outlet in Huntington, Vermont (Fig 9; Larsen, 1972, 1987; Larsen et al., 2003b; Wright, 2018a). The rapid drop in lake level associated with this breakout flood allowed much coarser sediment (derived from shore) to reach the site of the Waterbury Reservoir section.

The beginning of Glacial Lake Winooski is harder to date because no good sections of varved lake sediments have been found in the Stevens Branch valley, the valley hosting the lake during its earliest years immediately north of the lake’s Williamstown Gulf outlet (Figs. 7, 14).

However, the Muzzy Brook, Montpelier, and Wrightsville Reservoir varve records provide the basis for a good estimate. While only the Wrightsville Reservoir section begins directly above till, the Muzzy Brook and Montpelier sections both begin with thick, ice-proximal varves and likely record deposition shortly after the ice sheet retreated north of these respective locations. From the varve correlation graph (Fig. 14) it’s clear that sediments begin accumulating first at the southernmost section (Muzzy Brook, green line), begin next in Montpelier (grey line), and last at the northernmost section (Wrightsville Reservoir, orange line). Varves thin over a span of 20 to 30 years as the ice sheet, the source of most of the sediment, retreats farther away (Fig. 14).

Ice margin retreat rates can be calculated between Muzzy Brook, Montpelier, and the Wrightsville Reservoir using the time interval between the start of sedimentation at each site and the distance between these sites (Fig. 15,A,B). The calculated retreat rates between Muzzy Brook and Montpelier (295 m/year) and Montpelier and the Wrightsville Reservoir (305 m/year) are almost identical. Furthermore, Ridge and others (2012), using many more sections of varved lake sediments, calculate an identical retreat rate of ~300 m/year in the upper Connecticut River valley, the reach of the valley due east of the graphed sections (Fig. 12).

If the ice retreated north from the outlet of Glacial Lake Winooski to the latitude of the Muzzy Brook section at 300 m/year it would have taken ~50 years to retreat that distance (Fig. 15, D). A similar calculation based on a more conservative retreat rate of 250 m/year increases the duration of that retreat to 60 years (Fig. 15, E). These calculations indicate that the ice margin retreated across the outlet of Glacial Lake Winooski 50 to 60 years before the ice retreated north of Muzzy Brook at North American Varve 6658 putting the beginning of Glacial Lake Winooski at approximately North American Varve 6600 which corresponds to about 14,220 yr BP (Fig. 14). Therefore, Glacial Lake Winooski lasted ~375 years, expanding both north and west of its outlet before the ice dam in the Winooski River valley retreated far enough west to uncover the Gillette Pond outlet allowing water levels to rapidly fall and the Williamstown Gulf outlet to be abandoned (Figs. 6-9, 14). These dates are consistent with Ridge’s (2016) 14.1 and 13.9 ka timelines across the Winooski drainage basin (Fig. 12).

Further confirmation of the timing of ice retreat across the region comes from exposure age dating. In a study focused on sampling an elevation transect from Mount Mansfield in the northern Green Mountains (western flank of Glacial Lake Winooski, Fig. 12), Corbett and others (2019) found that 15 samples collected between ~400 and 1,200 m a.s.l. have indistinguishable 10Be ages averaging 13.9±0.6 ka indicating rapid thinning of the ice sheet at this time. This date range of rapid ice sheet thinning coincides with the ~14,170 to 13,800 ka life span of Glacial Lake Winooski when the growing lake tracked the rapidly retreating ice sheet east of the mountains. Exposure ages from elevation transects across New England similarly show 800-1,000 m of ice thinning over time spans too short to distinguish with this dating technique (Halsted et al., 2022).

The Littleton-Bethlehem Readvance of the ice sheet is well documented in northern New Hampshire by a widespread array of moraines (the White Mountain Moraine System) and a stratigraphy of till with thrust-faulted sand and gravel within Glacial Lake Hitchcock sediments near the Comerford Dam in the upper Connecticut River valley (Thompson et al., 2017; Thompson et al., 1999). Thompson and others (2017) associate this readvance with a 123 year long interval in the North American Varve Chronology (14,021-13,898 cal yr BP) that corresponds with the Older Dryas cooling event. This same readvance is also referred to as the Middlesex Readvance and appears as the 13.9 ka timeline in Figure 12 (Ridge, 2016; Ridge et al., 2012).

In the Winooski River basin many sites contain one or more layers of diamict interlayered with lake sediments and these have been cited as evidence of ice sheet readvance corresponding to the Littleton-Bethlehem (Middlesex) readvance (Dunn and Springston, 2019; Larsen et al., 2003a; Wright, 2015a). However, in many of these sections there’s no stratigraphic evidence recording ice advance or retreat below or above these diamict layers and landslides from the surrounding steep hillsides may be a more likely explanation for the diamict layers (Wright, 2023). Of particular note is a diamict overlying intensely sheared Glacial Lake Winooski sediments along Culver Brook adjacent to the Wrightsville Reservoir (Fig. 16; Larsen, 1999a; Stop 5C). Larsen (2001) dated a piece of wood found in the overlying till at 11,900 ±50 C-14 yr BP which corresponds to a calibrated age range of 13,675-13,791 yr BP, well after Glacial Lake Winooski had drained (Fig. 8). For this diamict to be a readvance till the ice sheet needed to advance and subsequently retreat across almost the entire Winooski River basin. While doing so it would have dammed the basin a second time, yet there’s no evidence of a younger widespread “Glacial Lake Winooski.” It seems more likely that the upper till at Culver Brook is a landslide deposit that occurred well after the ice had retreated from the area and climate had warmed sufficiently for the surrounding hill slopes to be tree-covered.

Our mapping elsewhere in northern Vermont indicates that landslides were common following ice retreat (Wright, 2023). A recent landslide (May, 2019) along Cotton Brook on the eastern slopes of the Green Mountains (~6.2 km SE of Stowe Village) reveals a stratigraphy consisting of multiple landslide deposits interbedded with undeformed lake sediments (Fig. 17). Some of the landslide deposits are diamict and others consist of faulted blocks of layered sand/silt that slumped into the lake basin. Note that the top several meters consists of diamict (colluvium) and before the landslide revealed the underlying lacustrine sediments, Wright (2018b) mapped this entire hill slope as till. Both recently deposited, unvegetated till and thick accumulations of lacustrine sediments are unstable on steep mountain slopes, particularly when water saturated. Whether triggered by storms, spring thaw, high groundwater pressure, or some other process, landslide deposits are common both within the region’s glacial lakes and on the steep slopes adjacent to these lakes. That said, many if not most of these deposits consist of diamict and distinguishing an in-place till vs remobilized till (diamict) is difficult or impossible in most areas. Detailed structural work designed to ascertain the direction of shear strain in deformed sediments underlying the diamict layers could help distinguish diamicts deposited by readvancing ice from those deposited by landslides. This work has not yet been attempted.

As the ice sheet thinned ice flow was directed parallel to the region’s valleys, the mountains emerged first as isolated nunataks and then as continuous ridges. We have mapped a variety of landform/sediment associations that likely formed along or close to the ice margin and provide some insight into processes occurring there.

In northern Vermont traditional recessional moraines, single or intersecting ridges largely composed of till, occur only rarely across valley bottoms and are small in scale. The near absence of this type of moraine is a likely consequence of the rapid ice sheet retreat rate across northern Vermont (~300 m/yr, see previous section). Prominent moraines, similar to White Mountain Moraine System associated with the Older Dryas readvance (Thompson et al., 2017; Thompson et al., 1999), have not yet been identified in the region.

While traditional moraines with positive relief are rare, another type of ice marginal landform is widely developed on many mountain slopes in the Green Mountains and may constitute a different style of moraine (Wright, 2019b). These are asymmetric steps or benches, the hill slope sequentially steepens and flattens, consisting almost entirely of till (diamict) without any positive relief (Figs. 18, 19). These landforms generally contour across the valley sides and commonly occur as flights of subparallel steps in areas where the till cover is thick enough to bury the underlying bedrock and on slopes range from ~12 to 20% (Figs. 18, 19). Steeper slopes are generally dominated by bedrock. These landforms are not well developed on gentle slopes <10% (Fig. 18). In some areas these landforms are cut by meltwater channels. Field work designed to understand the internal structure of these landforms has generally been unsuccessful. However, thin beds of silt/very fine sand occur in small exposures within the diamict, perhaps deposited in ice-marginal ponds, and these beds are universally deformed and generally dip gently upslope.

Wright (2019b, 2022d) proposed that these steps form when till, recently exposed as the ice sheet thins during the summer months, (1) flows down-slope and accumulates at the ice sheet margin and/or (2) is pushed into a ridge as the ice sheet expands upslope during the winter months. The thin layers of silt/fine sand may be deposited in small pools of meltwater collecting at the ice sheet margin. The upslope dip of these layers argues for (1) a limited degree of push from expanding ice or (2) that deformation may result from water-saturated till being squeezed from beneath the ice sheet margin. Till extrusion to the ice sheet margin and crevasses may also be a significant process and may explain some of the irregular patterns displayed by landforms. Davis and others (2022) refer to similar landforms they mapped on Mount Moosilauke as “till benches” and emphasize the role of till slumping against the ice sheet margin as key to their formation (Davis, 2024). While Geier (2021) suggested that these benches were produced by solifluction, Davis (2024) argues that these benches lack the distinctive pattern of solifluction lobes. We have similarly not observed any benches in the Green Mountains displaying features characteristic of solifluction lobes.

Where traditional recessional moraines record periods of time when till has accumulated at the ice margin, e.g. during a standstill or an advance, these benches/moraines, where well preserved, may record yearly thinning of the ice sheet. Along the flanks of Mount Mansfield and Belvidere Mountain Wright (2019b) measured vertical drops of 9-16 m between adjacent benches/moraines. Near the Québec border flights of meltwater channels drop 3 to 10 m between adjacent channels and may also record yearly ice sheet thinning (Wright, 2024). If these elevation drops are a measure of yearly ice sheet thinning, then 1,000 m of the adjacent mountains could have been exposed in just a few hundred years. While that rate seems very fast, it may be consistent with the calculated average ice retreat rates of ~300 m/yr and explain why exposure age dates show little change with elevation across New England (Corbett et al., 2019; Halsted et al., 2022).

Rapid thinning and retreat of the ice sheet across northern New England generated tremendous volumes of meltwater that flowed both beneath and along the margins of the ice sheet. Ice-marginal channels document the pathways this meltwater followed. A well-preserved set of meltwater channels occurs in the southwest corner of the Waterbury Quadrangle (Fig. 21; Wright et al., 2023a). These south- and west-sloping channels are parallel to the inferred margin of the ice sheet. These channels occur in sub-parallel sets and were likely sequentially eroded into recently exposed till by streams flowing parallel to the edge of the ice sheet as the ice sheet thinned from year to year (Fig. 20).

In addition to eroding channels, that same meltwater transported large volumes of sediment. Most of this sediment was eroded from till occurring beneath, on the surface of, or along the edge of the ice sheet, but also included debris entrained within the ice that was released as friction from moving water melted the adjacent ice. Kames, deposited by streams flowing between the edge of the ice sheet and adjacent mountain slopes, are particularly well developed along the western slopes of the Green Mountains, i.e. the eastern side of the Champlain valley, perhaps due to the large scale of the glacier in that valley (Connally and Calkin, 1972). The western extent of our recent mapping just reaches into the Champlain valley and this is where these landforms are best developed, albeit on a much smaller scale than at lower elevations farther west (Fig. 21). During times when the ice sheet bordered and/or dammed a glacial lake, sediments carried by meltwater streams formed ice-contact deltas where they entered that lake (Fig. 21). Several of these deltas were mapped where ice-marginal streams entered relatively small, short-lived glacial lakes dammed by the retreating ice along the western slopes of the Green Mountains (Wright, 2022c, d). Distributary channels are often preserved on the delta surfaces. Kettles are also common in both kames and ice- contact deltas (Fig. 21).

Eskers are common in the region and generally follow the major stream valleys either NW-SE across the mountains or N-S parallel to the mountains. This is consistent with the directions of ice sheet flow and subsequent hydraulic gradients as the ice sheet thinned and its flow directions were dictated by the region’s valleys. Ice flow across the mountains is highlighted by he Miller Brook esker which will be visited on this field trip (Stop B4). Meltwater flowing through this esker tunnel was sourced from ice in the Champlain valley, routed through Nebraska Notch, a low point in the Green Mountain ridge line, and then down the Miller Brook valley.

The Winooski River valley, prominently cutting WNW-ESE across the mountains, also hosted an esker. While a distinct esker landform no longer exists, its existence is indicated by numerous exposures of subaqueous fan deposits sourced from this esker, sedimentary structures preserved in a gravel pit in Middlesex, and numerous water well logs in the valley that bottom in coarse sand and gravel, i.e. the esker is buried beneath younger lacustrine sediments.

One of the best examples of an esker oriented N-S is a long segmented esker that follows Vermont Route 14 completely across the Montpelier Sheet and probably to the Québec border. This esker threads its way along numerous aligned stream valleys and up and over many low-elevation drainage divides (Fig. 22). While many portions of this esker have been quarried away, significant portions still exist and will be visited during the Friday field trip. The significance of this esker is that it demonstrates the evolution of a long-lasting, time-transgressive subglacial drainage conduit within the ice sheet during ice retreat. Given that eskers are time-transgressive depositional landforms, i.e. at any one time net accumulation of sediments likely occurred only in the last several kilometers of the subglacial ice tunnel, the up-glacier tunnel system itself must have continued to migrate northward as the locus of deposition in the subglacial tunnel also shifted to the north. In areas where no esker was mapped, it’s unclear whether (1) an esker was never deposited at these locations, (2) an esker was deposited but was later eroded away or fully excavated by human activities, or (3) the esker lies buried beneath younger, lacustrine deposits.

Alluvial fans occur on a wide-range of scales across northern Vermont. These form where sediments eroded from the valley sides, generally glacial till, have been carried downhill by tributary streams and deposited where the stream gradient abruptly lessens where it flows onto the valley bottom. The apex of these fans frequently consists of debris flow deposits. Farther down-fan sediments consist of lenses of sand/gravel that may fine to silt at their toes. In areas above the elevation of glacial lakes, the fans are deposited on till. More commonly they were deposited on lake bottom sediments and both concurrent with the lakes and the much longer period of time following. In many places these fans are large enough to extend almost completely across the valley and the courses of the local streams flow around the toes of these fans. Work on alluvial fans in northern Vermont suggests that fans have been episodically active throughout the Holocene and many received their most recent pulse of sediment following European land clearing in the late 18th and early 19th centuries (Bierman et al., 1997; Jennings et al., 2003). Related work by Noren et al. (2002) recording pulses of clastic sediment deposited in ponds and small lakes, indicates that pre-European settlement erosion has not been uniformly distributed throughout the Holocene and seems instead to be concentrated during periods of increased high-intensity storms. If climate shifts produce a greater frequency of high-intensity storms, further sedimentation on the area’s alluvial fans seems likely.

(Laurie Grigg)

Sediment cores from upland lakes and wetlands in central Vermont preserve an archive of post-glacial environments and provide insights on past changes in climate and vegetation. This presentation focuses on the transition from the late-Pleistocene to early Holocene in three records from the region, Twin Ponds, Brookfield, VT (Grigg et al., 2021; Mandl et al., 2016), Pecks Pond, Barre, VT (Magdon et al., 2018), and Knob Hill Pond, Marshfield, VT (Oswald and Foster, 2011; Oswald et al., 2018). The oldest radiocarbon ages from these records range from 13.0 to 13.7 cal ka BP but sediment recovery extends another 40-80cm suggesting these are minimum ages. Rapid warming centered around 11.5 cal ka BP, as well as, pre- and post-Younger Dryas centennial-scale climatic variability are prominent in all records. Multi-proxy records from Pecks Pond indicate an early warming (>13.3 cal ka BP) evident by an increase in CaCO3, aquatic vegetation, and bivalves. These results suggest that vegetation flourished in aquatic environments prior to terrestrial environments. Several centennial-scale, lake-level fluctuations at Pecks Pond are also evident >13.3 cal ka BP and may indicate variable precipitation that may have impacted the glacial mass balance of the retreating Laurentide Ice Sheet. At the start of the Younger Dryas cooling event (YD; ca. 12.9 cal ka BP), Pecks Pond fluctuated from shallow lake to wetland and back to shallow lake. During the YD, all proxies from Twin Ponds show evidence for generally cold but not necessarily reversing conditions and increased meltwater runoff. The Twin Ponds ostracod record shows a northward shift in the position of the summer polar front during the latter half of the YD which is consistent with evidence for increasing sea-ice instability in the northwestern North Atlantic. There is limited evidence for additional cool intervals centered around 10.6, 10.2 and 9.6 cal ka BP at Twin Ponds that may reflect continued instability in the North Atlantic during the final phases of deglaciation.

Pollen records from Twin Ponds (Grigg et al., 2021) and Knob Hill Pond (Oswald and Foster, 2011) suggest that central Vermont did not experience a significant tundra phase. Very low pollen accumulation rates prior to the YD suggest limited sedge, grass, and spruce or the long-distance transport of pollen grains from southern regions. Sedge pollen at the base of the pollen record from Twin Ponds indicates the establishment of a fringing wetland around the pond or limited patches of sedge tundra. During the YD, pollen accumulation rates of spruce increased slightly but herbaceous taxa remained low suggesting open or isolated patches of spruce amongst a largely till- dominated landscape. It was not until 11.5 cal ka BP at the end of the YD that pollen accumulation rates increase to levels comparable to a closed forest. This transition includes a sequences of non-analogue mixed forest types that culminated in the pine-oak forest characteristic of the early Holocene (Grigg et al., 2023).

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