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Few data are available to infer the thinning rate of the Laurentide Ice Sheet (LIS) through the last deglaciation, despite its importance for constraining past ice sheet response to climate warming. We measured 31 cosmogenic 10 Be exposure ages in samples collected on coastal mountainsides in Acadia National Park and from the slightly inland Pineo Ridge moraine complex, a ∼100-km-long glaciomarine delta, to constrain the timing and rate of LIS thinning and subsequent retreat in coastal Maine. Samples collected along vertical transects in Acadia National Park have indistinguishable exposure ages over a 300 m range of elevation, suggesting that rapid, century-scale thinning occurred at 15.2 ± 0.7 ka, similar to the timing of abrupt thinning inferred from cosmogenic exposure ages at Mt. Katahdin in central Maine (Davis et al., 2015). This rapid ice sheet surface lowering, which likely occurred during the latter part of the cold Heinrich Stadial 1 event (19–14.6 ka), may have been due to enhanced ice-shelf melt and calving in the Gulf of Maine, perhaps related to regional oceanic warming associated with a weakened Atlantic Meridional Overturning Circulation at this time. The ice margin subsequently stabilized at the Pineo Ridge moraine complex until 14.5 ± 0.7 ka, near the onset of Bolling Interstadial warming. Our 10 Be ages are substantially younger than marine radiocarbon constraints on LIS retreat in the coastal lowlands, suggesting that the deglacial marine reservoir effect in this area was ∼1,200 14 C years, perhaps also related to the sluggish Atlantic Meridional Overturning Circulation during Heinrich Stadial 1.

Fluvial, glacial, and estuarine deposits in the Delaware Valley record the response of the Delaware River to glaciation, sea-level change, and glacioisostasy during the Quaternary. Incision following an early Pleistocene glaciation created the present valley, which is inset into a Pliocene strath and fluvial plain. Middle and upper Pleistocene and Holocene deposits were laid down in this inset valley. Estuarine terraces in the lower valley and bayshore at + 20 m (probably Marine Isotope Stage [MIS] 11), + 8 m (MIS 5e), and + 3 m (MIS 5a or c), and a fluvial deposit that correlates to offshore MIS 3 marine deposits at − 20 m are at elevations consistent with glacioisostatic models. Successive incisions during lowstands in the middle and late Pleistocene lengthened, deepened, and narrowed the channel in the lower valley and shifted the channel westward in Delaware Bay. During MIS 2 glaciation, from 25 to 18 ka, the Delaware was diverted to the Hudson Shelf Valley by glacioisostatic tilting. Most glacial sediment was trapped in fluvial-lacustrine valley fills north of the terminal moraine. Incision of the valley fill was accomplished during the early stage of rebound, between 17 and 12 ka. Drainage to the Delaware shelf was restored between 15 and 13 ka as the forebulge collapsed. During incision, multiple postglacial terraces formed where the valley was perpendicular to rebound contours and so was steepened and elevated northward; and a single terrace formed where the valley paralleled the contours, and there was no differential elevation or steepening. About 65% of the original volume of MIS 2 glacial sediment remains in the main valley, and most of the eroded volume is in the channel in the lower valley beneath Holocene estuarine fill. Little glacial sediment reached the Delaware or Hudson shelf. Overbank deposition on the lower postglacial terrace and modern floodplain spans the Holocene. The volume of Holocene sediment in the estuary and bay yields a basinwide denudation rate of about 20 m/my.

Glacial retreat from valleys in the moderate relief (300–500 m) Small Lakes Section of the Appalachian Plateau in northeastern Pennsylvania was characterized by episodic deposition of till in a series of knobs that formed “beaded valleys”. Individual valleys have a north to south series of till knobs alternating with wetlands or lakes at a spacing of one to five kilometers. Outcrop and well data, while small and few in any individual knob, when put together from the more than 1000 knobs mapped in region, show that the till knobs are typically 30 to 50 m thick. The knobs are cored by subglacial till (lodgment-deformation till) with a wedge of supraglacial till (flow or re-sedimented till) on the south sides, push structures in the interiors and north sides, and an overall veneer of “colluviated till” that thickens downslope on all sides. Glaciofluvial deposits are scarce, usually appearing as thin lenses on the flank of the knobs. The knobs are interpreted to be the periglacially and post-glacially modified remnants of recessional moraines. Individual till knobs were rapidly deposited in a few decades, probably through layer by layer stacking of deformation till and till block melanges. Active ice shearing over inactive ice could form an adverse slope where rapid till deposition could take place. The ice retreated systematically in a stagnation-zone retreat mode, with active ice leaving till knobs and 1–5 km wide stagnant-zone ice leaving the lake basins between the knobs. The till knobs can be connected from valley to valley, in lines perpendicular to the southwesterly ice flow, to delineate ice margin positions across the region. Valleys transverse to ice flow have “till shadows”, 30 to 50 m thick till deposits on the north or lee side of the valley. “One-sided” post-glacial bedrock gorges with bedrock on the south side and till on the north side, are ubiquitous in “till shadow” valleys and form as the stream incises down the bedrock-till contact. Till outcrops in such gorges show no evidence of weathered older till under the late Wisconsinan till. The till is predominantly lodgment-deformation till in the lower two-thirds of the section and mainly ablation and meltout till above that. On the adverse slope in the transverse valley the glacier did not deposit till, possibly because it was on a “till starved” relatively steep bedrock slope rather than a “till rich” deformable bed of a till knob.

Publisher Summary There is evidence for at least four Pleistocene glacial advances into Pennsylvania. Each time the Laurentide ice sheet advanced, it entered Pennsylvania from the North-West and the North-East. The triangular shaped unglaciated area between the two lobes is called the Salamanca Re-entrant. In Northwestern Pennsylvania, the relatively low relief, shaly bedrock, and abundant debris from the Great Lake basins produced a dominance of deposition over erosion in each glacial advance. This produced multiple till-sheet sequences whose layers sometimes are separated by partly-preserved weathering profiles. These till sequences are well exposed in open pit coal mines. The chapter discusses that the pre-glacial stream drainage in north-western Pennsylvania was to the north-west. During glaciation this drainage was blocked and diverted to the south-west to form an ice marginal drainage system, the present Allegheny-Ohio River system. In north-eastern Pennsylvania the moderate relief on sandstone bedrock produced a dominance of erosion over deposition during each glacial advance. Each successive advance almost entirely removed the deposits of the previous advance and some bedrock. Other proglacial lakes were imponded along that limit, along younger limits, and North of the Late Pleistocene limit as ice receded from Pennsylvania.

From New Jersey to southwestern New York, Altonian colluvium and fluvial gravel have been recognized beneath Woodfordian drift, but the existence of Altonian till near or beyond the Woodfordian border has not been demonstrated. Reworking of weathered materials by Woodfordian ice has produced some drift with an “Altonian” appearance. Soil criteria, applied without regard for parent material differences, have been incorrectly used to infer an Altonian age for Illinoian deposits that have a redbed provenance. The currently mapped distribution of Altonian till and the occurrence of undisturbed Illinoian deposits within areas of inferred Altonian ice cover are incompatible with the behavior of ice sheets.

Abrahams' comment relates meander length to channel cross-section shape and recurrence interval so that meander length can be both directly and inversely proportional to rock resistance. This reply notes that either meander length is directly proportional to rock resistance or it is not; one cannot have it both ways. Many Appalachian Valley and Ridge bedrock meanders are shown to be the same size as alluvial meanders, and appear to be somewhat underfit. A hypothesis is proposed where modest discharge increases may have accelerated bedrock meander cutting although present streams remain capable of slowly cutting the meanders.

The marine 18 O record indicates that the first ice advance into the Appalachians equivalent to that of the late Wisconsin occurred in the Pliocene at about 2.4 Ma. Only periodic localized glacial and periglacial activity occurred in the northern Appalachians from 2.33 to 0.85 Ma. After 0.85 Ma (middle to late Pleistocene) eight out of ten glaciations covered the northern Appalachians and brought periglacial conditions to the southern Appalachians. During the last 850,000 years, glaciation extended to the late Wisconsin limit within the Appalachians for 15% of the time and extended to the early Wisconsin limit just north of the Appalachians for 25% of the time. Pleistocene glacial erosion of the Appalachians of 120–200 m is necessary to account for the estimated volume of Pleistocene ocean sediment derived from the Appalachians and the estimated volume of Pliocene-Pleistocene glaciogenic ocean sediment derived from the entire Laurentide ice sheet. This amount of glacial erosion, occurring over the maximum expectable duration of Pliocene-Pleistocene ice cover of 247,000 years, yields glacial erosion rates from 0.5–0.8 m ka −1 . Such depths of erosion by thick, warm-based ice streams were localized in the lowlands and estuaries of the northern Appalachians. In the Maritimes the presence of thin, cold-based ice caused minimal erosion of uplands. Periglacial landforms and tundra-boreal forest were distributed throughout the Appalachian Highlands and the northern Coastal Plain during the glacial maxima. Periglacial erosion of 0.15–0.30 m ka −1 near the ice in Pennsylvania decline to 0.1–0.2 m ka −1 in the southern Piedmont. Each cold period almost completely renewed the regolith by eroding slopes and filling valleys. The cumulative effect of the eight to ten major cold events could have produced as much as a few ten's of meters of periglacial ridge-top lowering during the Pleistocene. Both estimated glacial and periglacial erosion rates greatly exceed estimated current fluvial erosion rates. Therefore glacial and periglacial processes should be considered the dominant processes in shaping the present Appalachian landscape. The northern Appalachians may be evolving towards a form where the landscape provides just the shape necessary to permit the least work conveyance of the ice across the region during glacial maxima. The southern Appalachians may be evolving towards a form where the landscape provides just the hillslope form necessary to transport the debris provided by periglacial conditions.

Marked differences in bedrock meander dimensions in the Appalachian Valley and Ridge province, at times in adjacent reaches of a single stream, are related to differences in relative erodability of the bedrock. Meanders cut in thick-bedded to massive lithologies, typically carbonates, are distinctly smaller overall than meanders cut into shaly lithologies. Other factors that could affect bedrock meander dimensions were considered and none appeared to offer any help in explaining the dimensional differences observed above. In plateau regions underlain by essentially horizontal strata, the meander form characteristic of one lithology may become superimposed upon another as incision progresses, producing anomalous relationships between bedrock meander dimensions and lithology. Empirical relations developed from meander geometry measurements and estimated bedrock erodability for 78 bedrock meander reaches, containing a total of 1089 individual meander loops, show that meanders cut in shaly lithologies (ML = 105 Q0.50f, where ML = meander length and Qf = most probable annual flood) are about twice the length of meanders cut in non-shaly lithologies (ML = 39.30.56f). The valley floor width of the meanders cut in shaly bedrock (VF = 28 Q0.43f) is two to three times wider than the valley floor width of the meanders cut in the more resistant non-shaly bedrock (VF = 8 Q0.43f). The mean and median meander length values for individual reaches typically differ by less than 10 per cent.