The Late Quaternary Construction of Cape Cod, Massachusetts: A Reconsideration of the W.M. Davis Model

Geological Society of America Special Paper 309 1996

The Late Quaternary Construction of Cape Cod, Massachusetts: A Reconsideration of the W. M. Davis Model ABSTRACT Like the W. M. Davis construction of Cape Cod published in 1896, this special paper suggests that the Cape was formed by glacial deposition during the late Pleistocene and by marine and aeolian processes during the Holocene. It differs, however, from the Davis model in several significant ways. For example, Davis proposed that the lower Cape extended 4 km east of its present location, and that only 4,000 yr were needed for the lower Cape to attain its present form. This study indicates that the glacial Cape extended as far as 7 km east of its present position, and that it took about 9,500 yr for the Cape to attain its present morphology. Davis also believed that the detritus eroded from the sea cliffs on the east side of the lower Cape was transported northward to form the Provincetown Hook. On the other hand, this book indicates that, as a result of the subaerial exposure of Georges Bank from 9,500 to 6,000 yr ago, littoral drift to the north was inhibited and sediment was transported southward to fill a depression at the Cape’s elbow. Since Georges Bank became submerged about 6,000 yr ago, however, littoral drift shifted partly to the north, leading to the construction of Provincetown Hook, a process that is still taking place today. The bulk of material (86%) eroded from the cliffs along the east side of the lower Cape during the last 6,000 yr was transported northward. Of the remainder, 7% was used in the construction of the beaches and offshore bars fronting the cliffs, and the remaining 7% was incorporated into the spits and barriers south of the cliffs. During the last 70 yr, the eastern Cape cliffs have been retreating at an average rate of 0.8 m a–1, a rate enhanced by a relative rise in sea level of about 2 to 3 mm a–1. former glaciated terrain possessed irregularities comparable to those existing today, (3) the amount of newly formed land is less on protected than on exposed shores, and (4) cliffs protected by marshes and bars retreated before the bars and marshes were constructed in front of them. In his reconstruction, Davis also assumed that any changes in sea level were minor, and what changes occurred aided in the expansion of Cape Cod. According to Davis, the reconstructed glacial and fluvio-glacial terrain was bounded on all sides by steep cliffs of moderate relief. The reconstruction of Cape Cod using the above principles led to the extension of less than 1 km of land in the bays, more than 1 km to the west side of the Cape, and about 4 km to the east side of the lower Cape. Davis estimated that approximately 3,000 to 4,000 yr were adequate to accomplish the coastal retreat of nearly 4 km on the east side of the Cape. Sediments

INTRODUCTION The present morphology of Cape Cod and vicinity in southeastern Massachusetts is the result of Wisconsinan glaciation of a fluvial terrain and the subsequent modification of this glaciated landscape, principally by marine and aeolian processes during the Holocene. In his geographic essay on the outline of Cape Cod, Davis (1896; see also Davis, 1954) reconstructed the original glacial shape of the Cape Cod peninsula by removing the Holocene Provincetown Hook, the NausetChatham-Monomoy spits and bars, and a few small bars near Wellfleet, and by excavating the tidal marshes that were filled during the Holocene near Wellfleet, along the Pamet River and elsewhere. Cape Cod’s former seaward extension was reconstructed assuming that (1) subaerial erosion has not affected significantly the glacial topography since it was created, (2) the

Uchupi, E., Giese, G. S., Aubrey, D. G., and Kim, D.-J., 1996, The Late Quaternary Construction of Cape Cod, Massachusetts: A Reconsideration of the W. M. Davis Model: Boulder, Colorado, Geological Society of America Special Paper 309.

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eroded from the northern end of the eastward-facing cliff were hypothesized to be transported northwestward by longshore currents to build Provincetown Hook. Concurrent with the growth of the hook and erosion of the glacial sediments to the southeast, the fulcrum of shore erosion/deposition migrated northward, leading to the construction of a new tangential spit to the north and seaward of the previous one. As these spits grew northwestward, they curved and fused together, forming Provincetown Hook. According to Davis, the present morphology of the west coast of the Cape was due to a combination of northwest gales and shore currents that transported the sediments southward to form the tombolos of South Truro and Wellfleet prior to the development of Provincetown Hook, and southwest gales and northerly sediment transport since the development of the spit. He also believed that some southerly sediment transport is still taking place in western Cape Cod Bay today, contributing to the southwest-trending Jeremy Point and Billingsgate Shoal. In his discourse on the geomorphology of Cape Cod, Davis discussed neither the development of the north and south shores of “upper” Cape Cod (the east-west segment of Cape Cod), nor the origin of the barrier islands at the southeastern tip of the Cape. He did not describe the processes responsible for the creation of glacial Cape Cod, and he lacked offshore and onshore subsurface data that would have helped him constrain his reconstruction of the original outline of the glacial Cape. The effect of change in sea level was considered only in passing and the roles that glacial loading and unloading played in the creation of Cape Cod were not considered as their effects either were not recognized or were not deemed significant. Lacking a means of reliably dating its sedimentary deposits, Davis was unable to construct a reliable chronology for the development of the region. This book applies this new information, coupled with a better understanding of nearshore marine processes, to update Davis’s model on the genesis of Cape Cod. REGIONAL SETTING Physiography The Cape Cod peninsula, in the shape of a bent arm, is located in southeastern Massachusetts. It is separated from the mainland by a northeast-trending topographic low (Monument and Scussett Valleys) along which Cape Cod Canal was excavated (Fig. 1). The peninsula’s north shore from the Cape Cod Canal to 70°15′W is characterized by both erosional and depositional features, but farther eastward to Wellfleet Harbor the shore is mainly erosional (Fig. 2). Along the seaward side of Wellfleet Harbor (Great Island–Jeremy Point–Billingsgate Shoal) is another zone of deposition that becomes one of erosion slightly north of 41°55′N. This erosional segment extends to the southern tip of Provincetown Hook, a constructional feature. From the eastern side of the peninsula on the Atlantic Ocean side to about 41°55′N (Coast Guard Beach, Eastham), the coastal zone is one of extensive erosion. Farther south, in

the area of Nauset Beach and Monomoy Island, the Cape’s shore again is one mainly of construction, characterized by frequent and large changes in shore configuration. The Cape’s south shore facing Nantucket Sound is a low-energy environment with most of the shore undergoing erosion, interspersed with areas of deposition. The topography of the Cape is dominated by features produced by glacial, fluvio-glacial, lacustrine-glacial, and aeolian processes. A belt of morainal topography that dominates the Elizabeth Islands can be traced to the southwestern heel of Cape Cod at Woods Hole (Fig. 1). From there the Buzzards Bay morainal ridge, having a maximum elevation of more than 60 m, extends northeastward toward the east end of Cape Cod Canal, where it is truncated by the rounded and hilly Sandwich morainal ridge along the south shore of Cape Cod Bay (Fig. 2) (Woodworth, 1934a, p. 7, 1934b, p. 238; Mather et al., 1942). As the ridge extends eastward it gradually diminishes in elevation, finally merging with the outwash plains in the eastern upper Cape. North of this ridge is a series of deltaic terrains associated with a proglacial lake that existed in late Wisconsinan in the present geographic position of Cape Cod Bay. Descending gradually southward from the foot of the morainal ridge on the upper Cape is a series of outwash plains displaying declivities of about 0.2° (Uchupi and Oldale, 1994). As a result of postglacial marine erosion, the gradient of the plains increases to about 1° as the plains terminate along the south shore of Cape Cod. The lower Cape (the north-south segment that lies northeastward) is dominated by plains that descend gradually westward in the direction of Cape Cod Bay and display reliefs in excess of 40 m in the bluffs along the east coast of Cape Cod (Fig. 2). The bluffs, which are approximately 30 km long, have slopes in excess of 30° and form a broad arc slightly convex to the east. Continuity of the cliff face is disrupted by a series of gaps caused by breaching of the cliff by east-west–trending valleys (Fisher, 1987). Wave erosion at the cliff toe and the slides produced by this undercutting have led to westward retreat of these Atlantic cliffs at rates of about 1 m a–1 during the last 10 yr (U.S. Army Corps of Engineers, 1969). According to Zeigler et al. (1964b), the retreat rate from 1887 to 1957–1958 was slightly slower, about 0.8 m a–1. Miller and Aubrey (1985), who compared retreat rates on varying time scales using different measurement technologies, determined that the rate of retreat for the period 1970–1974 was 0.8 m a–1 and they estimated that the mean recession rate from 1879–1974 was 0.92 ma–1, a value comparable to those of

Figure 1. Bathymetry of Cape Cod and vicinity. From Uchupi (1987). TT = site of Texas Tower borehole; VC = site of vibracores. Also shown are the locations of seismic reflection profiles discussed in the text. Numbers beside locations indicate figure number of profiles. Inset shows location of Cape Cod and the topography of the southern New England offshore area. EI = Elizabeth Islands; JP = Jeremey Point; MV = Monument Valley and Scussett Valley; SP = Squibnocket Point; WB = Wilkinson Basin; WH = Woods Hole.

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Marindin (1889) and Davis (1896). The glacial outwash plain deposits on the western end of the plains in Truro in the more protected waters of Cape Cod Bay also are retreating, but at the much slower rate of about 0.3 ma–1 (Giese, 1964). Scattered throughout the outwash plains are numerous ket-

tle lakes ranging in diameter from several hundred meters to as much as several kilometers, with depths of several meters to more than 25 m (Oldale, 1992, p. 66). Water levels in the kettle lakes mirror groundwater levels in surrounding deposits. The lakes are particularly abundant in the Mashpee Pitted plain, the

Figure 2. Surficial geology and distribution of relict spring sapping valleys of Cape Cod. Geology from Oldale and Barlow (1986); valley distribution compiled from U.S. Geological Survey 7.5′, 1:24,000-scale quadrangle maps of the region are from Uchupi and Oldale (1994). Also shown are the locations of the geologic cross sections in Figure 6 and the positions of the boreholes in Figures 7 through 10. Postglacial deposits include fluvial, aeolian, marine, and beach sediments. BB = Ballston Beach; BH = Barnstable Harbor; BP = Beach Point (aka Pilgrim Beach); BR = Boat Meadow River; CH = Chatham Harbor; CO = Coonamesset and East Falmouth; GP = Gunning Point, Hamlin Point, Flume Pond; HH = High Head; LoP = Long Point; LP = Little Pleasant Pond; MH = Megansett Harbor; MI = Morris Island; NB = Nauset Beach; NH = Nauset Harbor; NMI = North Monomoy Island; OAF = Otis Air Force Base (now Massachusetts Military Reservation); OP = Oyster Pond; P = Pleasant Bay; PH = Phinney Harbor; POH = Pocasset Harbor; PPB = Popponesset Beach; QH = Quissett Harbor; RBH = Red Brook Harbor; SH = Squeteague Harbor; SMI = South Monomoy Island; SP = Salt Pond; WB = Waquoit Bay; WFH = West Falmouth Harbor.

Late Quaternary construction of Cape Cod, Massachusetts westernmost outwash plain on the upper Cape. Some of these topographic lows form coastal ponds separated from the open sea by baymouth bars. A few of them have natural connections to the open sea (i.e., Waquoit Bay, perhaps Salt Pond in Eastham, and Wellfleet Harbor), whereas others have been connected to the open sea by manmade channels (Oldale, 1992, p. 128). The surface of the plains is incised by valleys. These erosional features are unusual in that they lack modern streams because the plain sediments are too permeable to encourage runoff. The valleys tend to be straight rather than dendritic and have flat floors that are mantled with sand and gravel, short tributaries, and amphitheater-like heads; they also have retained their widths for considerable distances. The continuity of many of them is disrupted by kettles. The valleys also tend to stop short of the upper parts of the outwash plains where meltwater sources were located (Oldale, 1992, p. 72). For all of these reasons, Uchupi and Oldale (1994) proposed that the valleys were eroded by groundwater seeps fed by proglacial lakes (their high hydrostatic heads led to the elevation of the water table) that existed at that time rather than by glacio-fluvial meltwater from the Wisconsinan glaciers as proposed by Wigglesworth (1934, p. 132), Woodworth (1934b, p. 262–263), Zeigler et al. (1964c), and Strahler (1966, p. 19). The valleys are most abundant in the Mashpee Pitted plain in the upper Cape, where they drain southward toward Nantucket Sound (Fig. 2). Their lower reaches are drowned to form narrow bays trending at an angle to the present shore. Some of the bays are located along trunk valleys, and others are superimposed on their tributaries. The valleys in the lower Cape plains, known as hallows or pamets, drain westward in the direction of Cape Cod Bay. One of them, Pamet River, has been eroded below present sea level and extends across the width of the lower Cape from Cape Cod Bay to the Atlantic Ocean (Fig. 2). Today, Pamet River is separated on the east from the open ocean by a narrow strip of beach (Ballston Beach) that is eroded and sometimes overwashed during intense northeast winter storms, although restored by renewed deposition within a few weeks after the storm (Godfrey and Leatherman, 1979). The upper Pamet is separated from Cape Cod Bay by a oneway valve in a drainage pipe under a local road approximately 100 m west of State Highway 6. At the elbow of Cape Cod are two major barrier beach systems, Nauset Beach and Monomoy Island (Fig. 1). West of Nauset Beach is a broad low occupied by marshes, Chatham and Nauset Harbors, and Pleasant and Little Pleasant Bays. Bordering the marshes are scarps that may be relict sea cliffs (R. N. Oldale, cited in Aubrey et al., 1982). Historically, the shoreline along the elbow of Cape Cod has been subject to frequent changes in shape and position. The northern tip of the Cape is dominated by Provincetown Hook made of a series of concentric spits capped by the most extensive dune field in the region. The hook terminates east-southeastward against the east-trending former sea cliffs of High Head, which rise nearly 20 m to the Truro outwash plain. These bluffs mark the northern edge of

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the exposed glacial deposits of the Cape. North of the cliffs are two small outliers of the Truro outwash plain that are blanketed by the dune sands of the Provincetown Hook (Fig. 2). Another spit is found along the north coast of the Cape, the east-trending Sandy Neck that protects Barnstable Harbor from Cape Cod Bay. Capping this spit are dunes second in size to those in the Provincetown Hook (Leatherman, 1988, p. 36). Adding to the morphology of the Cape’s landscape are the numerous salt marshes behind barrier beaches and spits, and along the upland shores of bays and margins of tidal creeks. Enclosed within the bent arm of the Cape is Cape Cod Bay whereas south of the Cape lie Nantucket and Vineyard Sounds and the Elizabeth Islands, Martha’s Vineyard, No Mans Land, and Muskeget, Tuckernuck, and Nantucket Islands (Fig. 1). The Elizabeth Islands are the mostly submerged extension of the southwesterly trending Buzzards Bay moraine. The latter five islands are the subaerial segments of the ice-sheet margin that extends from Long Island (New York), to the Grand Banks of Newfoundland via Block Island (Rhode Island), Georges Bank, and the offshore banks on the Nova Scotian Shelf. The islands are separated from the mainland and Cape Cod by Vineyard and Nantucket Sounds that have average depths of slightly more than 10 m. Scattered throughout these sounds are shoals displaying reliefs of about 5 m and aligned in a complex fashion with the tidal currents of the region. Atop these shoals are smaller sediment waves generally at right angles to tidal flow. Southeast and east of Nantucket Island is an extensive area of shallow topography known as Nantucket Shoals. The area is dominated by northeast-trending shoals with reliefs in excess of 10 m parallel to the tidal currents (Uchupi, 1968). Along the crests of the larger shoals are linear sediment ridges also aligned parallel to the tidal currents; superimposed on the crests of the shoals and the troughs between them are smaller waves, most of which are at right angles to the flood and ebb tidal currents. As a result of its shifting shoals, some of which are less than 1 m below sea level, Nantucket Shoals is considered one of the greatest hazards to navigation in the eastern United States. Southwest of Nantucket Shoals is the east coast continental shelf whose smooth surface is broken by erosional features such as channels and terraces and depositional features such as sand swells (Uchupi, 1968). A broad amphitheater shape low on the outer shelf known as the mud patch has an extensive fine-grained sediment cover (Fig. 1). North of 41°20′N, Nantucket Shoals change their orientation from northeast-southwest to north-south as they adjust themselves to a change in the tidal current direction. East of these shoals the bottom contours bulge eastward in the form of a fan to which the name Fan A is suggested (Fig. 1). Offshore from Monomoy Island and Nauset Beach, the sea floor is dominated by northeast-trending swells that can be traced to a depth of about 20 m. Beyond that depth, sea-floor roughness diminishes to the near 40-m depth, and sea-floor declivity increases as it descends to the Wilkinson Basin complex in the Gulf of Maine (Fig. 1). From 41°45′N to 41°50′N, the surface of the

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inner shelf shallower than 10 m is characterized by a ridgetrough topography paralleling the shore with a relief of several meters (Aubrey et al., 1982). The outer shelf to a depth of about 40 m and the slope beyond that depth are relatively featureless. This slope segment between Fan A and Fan C is in the form of an embayment in the general topographic trend, reflecting a now-filled depression seaward of Nauset Beach. Although the construction of Fan B on the upper slope by the sediments filling the low has led to eastward progradation, the slope in the low is still not aligned with the slope north and south of the depression. The shelf north of 41°50′N to the northern tip of Cape Cod displays the same morphologic features as the shelf segment farther south, an inner ridge–trough zone in depths of water less than 10 m, and a smooth outer shelf to the break in slope at a depth of about 40 m. The slope seaward of the shelf’s edge consists of a fan (Fan C) whose topographic expression can be traced to a depth of 180 m (Fig. 1). The major topographic features in Cape Cod Bay are Fishing Ledge and Billingsgate Shoal (Fig. 1). The elliptical-shaped Fishing Ledge in the center of the bay has about 6 m of relief, and the 8-m-high Billingsgate Shoal aligned obliquely to the west coast of Cape Cod extends southwestward for a distance of about 10 km. East of Billingsgate Shoal, just south of Jeremy Point, is Billingsgate Bank, marking the former position of the island of that name which submerged in 1942 (Leatherman, 1988, p. 82). The sea floor along the western periphery of Cape Cod Bay in depths of less than 10 m, where basement is exposed, is quite rough with numerous small highs rising about 2 m above the surrounding sea floor. The sea floor north of Cape Cod Bay is dominated by the flat-topped Stellwagen Bank that mimics the general configuration of the lower Cape. This topographic high with depths of less than 40 m along its top is separated from the northern tip of Cape Cod by Little Stellwagen Basin. West of the bank is the deeper Stellwagen Basin (Fig. 1). The sea floor west of this depression, at depths of less than 40 m where basement is exposed or thinly mantled by sediments, is rough with numerous topographic irregularities displaying reliefs of several meters. North-northeast of Stellwagen Bank is a plateau-like feature broken into a series of highs and lows dominated by Tillies Basin (Fig. 1). East of Cape Cod is the Gulf of Maine, a rectangular depression with an average depth of 150 m consisting of a series of northeastand east-trending basins separated by swells, ridges, and flattopped banks (inset map, Fig. 1). Many of these basins have depths in excess of 200 m, although one of the shoals is just 9 m in depth. The Gulf of Maine depression is separated from the open sea by Georges Bank with two channels, Great South Channel east of Nantucket Shoals (Fig. 1) and Northeast Channel east of 67°W, that serve as deep-water passageways between the Gulf and the open sea. Geologic setting Nantucket Sound and Islands. The basement bedrock underlying the region consists of Precambrian and Paleozoic

rocks and Late Triassic to Early Jurassic rift system rocks (Emery and Uchupi, 1984; see also references therein). The pre–Late Triassic rocks reflect tectonism associated with opening and closing of a pre-Mesozoic Atlantic Ocean, and the Late Triassic–Early Jurassic structures reflect the breakup of Pangaea and the formation of the present Atlantic. An unconformity-truncating basement formed from uplift associated with the final continental decoupling and initiation of sea-floor spreading in the Atlantic in earliest Middle Jurassic (Uchupi and Emery, 1991, and references therein). Basement rocks are exposed only on the Massachusetts mainland and immediately offshore; on the Cape and the Islands, the basement is buried by coastal plain and/or glacial deposits. Martha’s Vineyard and Nantucket Islands (offshore islands south of peninsular Cape Cod) consist mainly of coastal plain deposits of Cretaceous and Cenozoic age capped by a thin veneer of Wisconsinan glacial sediments in the form of outwash plains and moraines (Oldale, 1982; 1992, p. 21). On Sankaty Head, Nantucket Island (Fig. 1), the Wisconsinan sediments rest on marine sediments having a uranium-thorium age of 133,000 ± 7,000 B.P. and an amino-acid racemization (AAR) age of about 120,000 to 140,000 B.P., suggesting a Sangamonian age (Stage 6e) for the deposits (Oldale et al., 1982; Oldale and Eskenasy, 1983; Oldale and Colman, 1992). Below the marine deposits is another till, probably of Illinoian age. This lower till, which extends from Maine to southern New England and Long Island, New York, indicates that the Illinoian glaciation (Stage 6?) was at least as extensive as the late Wisconsinan Laurentide ice sheet (Oldale and Colman, 1992). Eroded on the Wisconsinan outwash deposits of Martha’s Vineyard and Nantucket Island is an extensive network of valleys that have flat floors and uniform slopes throughout their lengths (Woodworth, 1934b; Wigglesworth, 1934). The coastal ponds of Martha’s Vineyard occupy the southern terminus of these valleys; in Nantucket the valleys are truncated by the present shore, with three of them having ponds aligned parallel to their axes (Oldale, 1992, p. 69). Uchupi and Oldale (1994) have proposed that this valley network was eroded by seepages from a Pleistocene-age lake that occupied the present position of Nantucket Sound. The relief of the islands is mainly the result of fluvial erosion during several marine regressions extending as far back as the Oligocene Epoch. During these regressions, several lowlands (Nantucket Sound, Gulf of Maine, and the basins on the inner Nova Scotian Shelf) and cuesta segments (Long, Block, Elizabeth, Martha’s Vineyard and Nantucket Islands, Georges Bank, and the banks on the outer Nova Scotian Shelf) were eroded out of the continental shelf sediments. Vineyard Sound, the channel between Martha’s Vineyard and Elizabeth Islands and western Cape Cod, and Great South and Northeast Channels on Georges Bank, represent water gaps through the cuesta (Oldale and Uchupi, 1970; Emery and Uchupi, 1984, and references therein). During the intervening transgressions the fluvial terrain was partially or completely buried, to be partially or completely exhumed dur-

Late Quaternary construction of Cape Cod, Massachusetts ing the following regression. These episodes of erosion and deposition account for the gaps in the coastal plain sedimentary record in the lowlands and cuesta. During the later Wisconsinan glaciation, southward overthrusting of the coastal plain strata and outwash sediments by an advancing ice-sheet margin accentuated the fluvial relief of the coastal plain cuesta in the area of Gay Head, Martha’s Vineyard. Seismic reflection profiles recorded in Nantucket Sound, the lowland north of the Martha’s Vineyard/Nantucket cuesta, indicate that basement in the region is mantled by coastal plain sediments. The surface of the basement displays evidence of fluvial erosion (Oldale et al., 1973) with the valleys draining southward away from an east-trending 50-m high beneath the upper Cape (Oldale, 1969). A geophysical reconnaissance of the sound by Oldale et al. (1973) suggested that the coastal plain strata pinched out some distance south of Cape Cod. A recent more detailed investigation by O’Hara and Oldale (1987) indicated, however, that the sediments may extend across the width of the Sound to at least and possibly beneath the south shore of Cape Cod. To date, there are no borehole data to verify the presence of coastal plain deposits beneath the south coastal zone of the upper Cape. The surface of the coastal plain sediments beneath Nantucket Sound is cut by an extensive system of V-shaped channels, some of which may extend through the sediments onto the basement beneath them. Similar valleys little affected by subsequent glacial processes have been recognized as far west as Long Island Sound (Lewis and Needell, 1987; Needell and Lewis, 1984; Needell et al., 1987). Oldale et al. (1973) proposed that the valleys in Nantucket Sound and the other sounds were remnants of the fluvial system that carved a cuesta (Long Island, Block Island, Elizabeth Islands, Martha’s Vineyard, Nantucket Island, Georges Bank) and lowlands (Long Island, Block Island, Rhode Island, Vineyard and Nantucket Sounds, Gulf of Maine) out of the continental shelf strata during one of the preglacial regressions. According to O’Hara (1981), the fluvial terrain that influenced the flow of the late Wisconsinan Laurentide ice sheet played a significant role in the formation and locations of the end moraines. For example, the moraines in Martha’s Vineyard and Nantucket Island were controlled by the valley heads as the ice advanced against the preglacial fluvial divide, and formation of the Buzzards Bay moraine during a glacial readvance was controlled by the cuesta slope. The geophysical investigation by O’Hara and Oldale (1987), however, does not appear to substantiate the theory of a fluvial origin for these valleys. Their study shows that the thalwegs of the valleys deepen toward the center of Nantucket Sound away from both Cape Cod and Martha’s Vineyard and Nantucket Island, but these lows are not connected and do not define a trunk river that drained the lowland. The pristine V-shape of the channels also is difficult to explain in a region that was so extensively glaciated during the Pleistocene. Fluvial channels cut into the erodible coastal plain sediments prior to the Pleistocene refrigeration should have a more rounded U-shape typical of a glaciated terrain. This absence of glacial

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modification and lack of a trunk system led Oldale (personal communication, 1993) to propose that the valleys in Nantucket Sound are not of fluvial origin, but are tunnel valleys cut by subglacial meltwater processes. A similar origin has been proposed by Boyd et al. (1988) for the subsurface channel networks on the outer Scotian Shelf. If the channels in Nantucket Sound, and possibly the other sounds, were eroded by subglacial processes, then the hydraulic gradient of the subglacial water must have been great enough to permit their extension up the back slope of the cuesta. Resting on the subglacial eroded surface of the coastal plain sediments are acoustic units characterized by flat-lying gently dipping reflectors that O’Hara and Oldale (1987) interpreted as outwash and ice-contact stratified drift; other units of acoustically well-layered sequence they interpreted as glaciolacustrine in origin. These sediments, which locally may be as thick as 160 m in the valleys cut in the coastal plain sediments, are exposed in many topographic lows in Nantucket Sound. Present-day marine erosion is producing or maintaining the exposures of the glacial deposits. The drift sequence is capped by a major unconformity that O’Hara and Oldale (1987) believed was fluvially cut in postglacial time during a low sealevel stand and locally by a planar surface inferred by them to have been eroded during the subsequent rise in sea level. The valleys in the fluvially carved surface appear to be southerly prolongations of the spring sapping ones in the upper Cape (Uchupi and Oldale, 1994). Above these local fluvial unconformities are sediment patches that were inferred by O’Hara and Oldale to be of fluvial, freshwater lake(?), estuarine, and marine origin. Nantucket Shoals. Vibracores recovered from the western edge of Nantucket Shoals near Nantucket Island indicate that the upper 2 m of the shoals in this region consist of a mixture of coarse-to-fine quartz sand containing some sparse glauconite clasts and shell fragments. At some of the core sites, the sands also contain appreciable amounts of silt and clay. Except at core SPC6 where the finer sediment occurs at the surface, the finer sands are found near the base of the sediment section sampled. The sands in vibracore SPC2 and SPC3 rest on a gravelly silty sand containing shell and barnacle fragments that we interpreted as glacial outwash with admixed recent sediments. On vibracore SSC3 the sands are underlain by a pebbly, black silty clay with a mild smell of H2S, probably a marsh deposit (Fig. 3). A boring drilled southeast of Nantucket at 40°58.4′N, 69°23.0′N at a depth of 18.5 m by J. M. Zeigler and associates at the Woods Hole Oceanographic Institution during the Texas Tower studies (Fig. 4) (Groot and Groot, 1964; Livingstone, 1964; Emery and Uchupi, 1972, p. 88, 93) on one of the shoals consisted of 27.7 m of fine sand atop a silt of unknown thickness (20 m was penetrated by the borehole) containing reworked Eocene spores and pollen, a few foraminifera, marine diatoms, sponge spicules, dinoflagellates, and several shells of Crepidula fornicata. One of the shells 1.5 m below the top of the silt (47.8 m

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Figure 3. Lithology of vibracores recovered near the east coast of Nantucket Island along the western edge of Nantucket Shoals. See Figure 1 for site of vibracores.

below sea level) yielded a radiocarbon date of 11,465 ± 400 B.P. (Groot and Groot, 1964). A boring a few meters from the above position drilled to a depth of 40 m below the sea floor yielded lenses of gravel-to-coarse sand and medium-to-fine silty sand. The sands in both of these borings, as with the sands at the vibracore sites, show evidence of reworking. A seismic reflection profile trending northeastward from the eastern edge of Nantucket Shoals into the Gulf of Maine displays two units. The lower unit has an irregular surface and displays discontinuous internal reflectors; we believe that this layer is of glacial origin, probably an ice contact facies. Draped over the lower unit and smoothing its surface irregularities are sediments which we inferred to be a postglacial marine sequence (Fan A: Fig. 5). Detritus making up this upper unit was derived from Nantucket Shoals to the west, as evidenced by its thickening in that direction. From the above data we infer that Nantucket Shoals have undergone and are still undergoing considerable reworking by

Figure 4. Lithology of wells drilled on Nantucket Shoals by J. M. Zeigler and associates during the Texas Towers investigation. This log was compiled by Emery and Uchupi (1972) from unpublished data archived at the Woods Hole Oceanographic Institution. See Figure 1 for location of borehole.

tidal currents. The ridge-trough morphology of the region is probably the result of the tidal current modification of an outwash plain deposited by the South Channel lobe to the east. Thus, the shoals may be marine equivalents to yardangs (streamlined erosional features aligned in the direction of the wind) on land. The larger shoals have probably remained stationary since their formation; only their surfaces are in constant change in response to the strong tidal currents. Tidal currents in this region are strong (up to 1 m s–1), although tidal vertical amplitudes are small. The tide here is an interaction of the MidAtlantic Bight tide and the Gulf of Maine tide. Differences in amplitude and phase of the two tides result in small ranges but strong currents, analogous to the tidal interaction in Nantucket Sound (Redfield, 1980, p. 82). The strong tides permit reworking and erosion of the sediments. Stratigraphic data from Texas Tower borings and vibracore

Late Quaternary construction of Cape Cod, Massachusetts

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Figure 5. Single-channel seismic reflection profile of Fan A east of Nantucket Shoals. B = basement; G = glacial sediments; H = Holocene postglacial marine sediments; M = acoustic multiple. Record is from the archives at the Woods Hole Oceanographic Institution and was recorded by E. Uchupi and R. D. Ballard during R/V Gosnold cruise 204 in 1973. Sound source: 40 in3 air gun with bubble suppressor. Profile 15 in Figure 1.

SSC3 suggest that some of the smaller features may have migrated some distance from their original positions, coming to rest over sediments indicative of less dynamic environments. The fines winnowed from Nantucket Shoals during the reworking of the sediments by tidal currents were transported not only east to Fan A, but southwestward to the Mud Patch, a topographic low on the outer shelf south of Nantucket Island (Fig. 1) (Bothner et al., 1981; Twichell et al., 1981). This amphitheater-shaped low that Uchupi (1967, 1970) believed was formed by the seaward gravitational sliding of the shelf strata is covered by an extensive blanket of muddy sediments (Schlee, 1973). Using coring, echo sounding, seismic reflection, side-scan sonar data, and 14C-age dates, Bothner et al. (1981) and Twichell et al. (1981) determined that the finegrained sediments in the Mud Patch rest on a relict undulating sandy surface formed during the Holocene transgression. They are as much as 13 m thick, were derived from both Nantucket Shoals and Georges Bank, and began to accumulate 10,950 to 7,850 B.P., when sea level was 45 to 15 m below present. According to Bothner et al. (1981) 210Pb inventories and tracemetal profiles indicate that the Mud Patch continues to be a sink of modern fine-grained sediments. The present rates of deposition of the fine-grained sediment being removed from Georges Bank and Nantucket Shoals range from 25 cm ka–1 in the center of the Patch, where the finest sediment occurs, and about 50 cm ka–1 in the eastern edge of the patch, where the sediments are sandier. Cape Cod. Basement beneath the upper Cape at a depth of 100 m to less than 50 m below sea level, forms an eastwest–trending swell divided by a south-trending valley extending from Cape Cod Bay to Nantucket Sound (Fig. 18) (Oldale,

1969). In the lower Cape, basement is at a depth of 125 to 150 m, disrupted by an east-trending Mesozoic rift at 42°N (Oldale, 1969; Ballard and Uchupi, 1975). Although glauconite of probable Miocene age has been noted by C. E. Reimers (in Oldale, 1976) in the drift of the northern and eastern upper Cape, no coastal plain sediments have been encountered by wells in the subsurface of the region (Fig. 6). Lack of velocity contrast between the glacial and nonglacial sediments also made it impossible to detect them, if present, during the geophysical survey of the Cape by Oldale (1969). The glacial sediments apparently rest directly on basement with the sequence immediately above basement consisting of poorly sorted till capped by fine sand, silt, and clay (Fig. 6). The finer grained sediments may represent a lacustrine facies that accumulated in a proglacial lake that extended from the northern upper Cape to the south across Nantucket Sound (Oldale, 1992, p. 75). Some segments of the region may have risen above the general level of the lake, as suggested by the absence of such strata on well MM drilled immediately to the north of the ice contact sediments near 70°30′N (Fig. 7). Resting on the lacustrine strata are outwash deposits of the Mashpee, Barnstable, and Harwich plains, a complex of subdeltas graded to the lake described above. These delta complexes were created by a system of braided streams debouching southward from the ice front to the north. At the distal ends of the outwash, plains were probably sandy delta fronts whose dips may have been comparable to the 10° to 25° gradients reported from sandy deltas elsewhere (Elliot, 1978). West of the Mashpee plain is the Buzzards Bay moraine, and north of the Mashpee and Barnstable outwash plains lies the Sandwich moraine (Fig. 2). These ridges were formed by the tectoniza-

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Figure 6. Geologic cross sections of the upper Cape. Compiled from LeBlanc et al. (1986, Sheet 1), Oldale and O’Hara (1990, Figs. 1, 8, 9), and O’Hara and Oldale (1987, Figs. 2 through 6). See Figure 2 for locations of sections.

tion of the Mashpee and Barnstable outwash deposits by overriding ice during periodic readvances of the retreating Buzzards Bay and Cape Cod glacial lobes (Oldale and O’Hara, 1984). The Harwich plain sediments with its ice contact head to the north (from the Cape Cod Bay lobe) and to the east (from the downwasting stagnate front of the South Channel glacial lobe) were not involved in this deformation, evidence that they are younger than the Sandwich moraine and the Barnstable and Mashpee plains. Coeval to the Harwich plain sediments are the ice-contact Nauset Heights deposits near the elbow of Cape Cod (Fig. 2) (Oldale, 1976). They are believed to have been deposited by the South Channel lobe in an interlobe between the Cape Cod Bay and South Channel lobes. North of the Sandwich moraine and the Harwich outwash plain is a lacustrine drift sequence deposited in a proglacial lake, Cape Cod Bay Glacial Lake, that existed at that time in the present geographic location of the bay. Additional evidence for the existence of this lake comes from outwash deltas of the

lower Cape and the lacustrine sediments along the west side of Cape Cod Bay (Oldale, 1976, 1982). This lake, which was dammed to the south by the Sandwich moraine and outwash plains of the upper Cape, by the Cape Cod Bay glacial lobe to the north, and by the South Channel lobe to the east, increased in size as the Cape Cod Bay lobe retreated northward. A pavement of ventifacts on the Wellfleet outwash plain in a sea cliff near Truro suggests that Cape Cod Bay Glacial Lake may have experienced at least one episode of lake level lowering, then rising before the final lowering lake stages (Oldale, 1976). Groundwater seepages from this lake carved the valleys in the permeable glacial drift of the upper Cape, damming the lake to the south. According to Oldale (1976; 1992, p. 73) the lake had its highest stands (25 m and 18 m above present sea level) early in its history and progressively lower stands as it grew eastward and northward where it found lower outflow routes. The lake’s earliest outflow apparently was via Monument valley, the location of present Cape Cod Canal, with later ones taking place via

Late Quaternary construction of Cape Cod, Massachusetts

Figure 7. Lithology of wells drilled on Mashpee Pitted (MM) and Eastham (6) plains. Compiled from LeBlanc et al. (1986, Sheet 1). See Figure 2 for locations of wells.

Bass River in the central upper Cape and Town Cove near the elbow of Cape Cod (Fig. 2). The shallow bays in the elbow of Cape Cod marked the position of a sublobe of the South Channel glacial lobe. When this sublobe retreated, its overlying sediments collapsed to form a topographic low, which was later filled to nearly sea level to form the present shallow bays. The lower Cape is dominated by westerly dipping outwash deposits indicating an easterly source in the South Channel lobe, and their texture indicates that this source was located some distance east of the present coast (except for the Eastham plain, which consists on the east of icecontact deposits). The oldest of these plains, the Wellfleet plain, is higher than the Highland and Truro plains to the north and Eastham to the south. The lower sequences in this plain (boreholes 2 through 5, Fig. 8) are fine deposits displaying foreset bedding suggestive of outwash delta sediments deposited in Cape Cod Bay Glacial Lake (Oldale, 1976). The upper sediments of the Wellfleet plain display a fabric characteristic of

11

Figure 8. Lithology of wells drilled on Wellfleet plain. Compiled from LeBlanc et al. (1986, Sheet 1). See Figure 2 for locations of wells and Figure 7 for lithology legend.

fluvial deposition. Diapirs of silt and clay that appear to be similar to the mudlumps on the Mississippi Delta (Morgan et al., 1968) intrude the Wellfleet plain deltaic sediments along the Atlantic sea cliff (Oldale and Barlow, 1986). Exposures of the Wellfleet and Highland sediments along the Atlantic coast sea cliffs have yielded cobbles of shelly marl of Eocene age (Crosby, 1879). Other reworked material from this and other plains in the Cape are arkosic sandstone fragments (abundant in the Nauset Heights deposits) resembling the Triassic rocks in the Connecticut Valley and the Bay of Fundy, Eocene spores and pollen, silicified wood fragments of probable Tertiary age, Tertiary(?) fish teeth, Tertiary and Pleistocene carbonized wood, and Pleistocene shells (Oldale, 1976). The fine-grained laminated clay and silt of the Highland plain, a small triangular-shaped area between the Wellfleet and Truro plains (Fig. 2), may represent lacustrine sediments deposited in a lake dammed by the Cape Cod Bay lobe to the north, the Wellfleet plain to the south, and the South Channel lobe to the east (Oldale, 1976). Seeps from this lake are respon-

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sible for the valleys eroded on the plains on the lower Cape. As these valleys also are cut into the Truro plain, the lake must have extended beyond this depositional unit. Uchupi and Oldale (1994) have proposed that the Pamet River (Fig. 2), which extends across the width of the lower Cape, may have breached the dam of the lake, allowing it to drain westward into Cape Cod Bay. They also inferred that this drainage was catastrophic, widening and deepening the Pamet River valley to below present sea level. The Truro plain defines the northern limit of outwash deposition in Cape Cod and is lower and younger than Wellfleet and Highland plains; it rests unconformably on the Highland plain sediments along the Atlantic sea cliffs (Oldale, 1976, 1979). The Truro sediments display large-scale foreset bedding in the cliffs along the Cape Cod Bay shore, evidence of a deltaic facies deposited in Cape Cod Bay Glacial Lake. The lower sediment sequence forming the Truro plain (boreholes 1 and D: Fig. 9) tend to be finer grained than the upper units. Those at borehole D contain traces of glauconite, shell fragments, carbonized wood, and Eocene spores and pollen. Zeigler et al. (1960, 1965) believed that the spores and pollen and carbonized wood in borehole D were in place and proposed that the fine-grained sediments in well D and the wells in the Provincetown Hook (boreholes A and C: Fig. 10) were Eocene in age. The presence of tests of Elphidium at a depth about 71 m, 14C dates of more than 42,000 yr for the carbonized wood, and sediment occurrence at the same depth as the well-stratified acoustic sequence in Cape Cod Bay that Hoskins and Knott (1961) inferred to be of Tertiary age confirmed the age of the sediments. In contrast, we believe that the spores and pollen were not deposited in place and that the fine-grained sediments deposited in Cape Cod Bay are glaciomarine in origin (Oldale, 1988). The basis for this conclusion includes the following evidence: the fine-grained sediments contain traces of glauconite, which is more characteristic of Miocene than Eocene sediments—thus they are younger than Eocene; the genus Elphidium is more typical of Plio-Pleistocene–Holocene than Eocene sediments; and, finally, extrapolation from coastal exposures indicates that the well-stratified sediments on Cape Cod are Wisconsinan lacustrine and not preglacial coastal plain sediments. The Eastham plain deposits, mainly of fluvial origin, are the youngest drift in Cape Cod. Prior to its creation and during the deposition of the other units, the Eastham plain, near the Cape’s elbow, was the site of a westward-trending glacial sublobe of the South Channel lobe. Like the other plains of the lower Cape, it, too, was deposited by the South Channel lobe, but graded to a lower Cape Cod Bay Glacial Lake level than the Truro plain, the next oldest drift. The surficial sediments in the Provincetown Hook west of the High Land cliff are mostly medium–to–very coarse quartzose sands with some rounded pebbles, carbonized wood, traces of glauconite, finer aeolian sands, and marsh deposits (boreholes A-C, AA, and BB: Fig. 10). The reddish clay in borehole BB near Pilgrim Lake was interpreted by Zeigler et al.

Figure 9. Lithology of wells drilled on Truro plain. Description of well D is based on the microscopic analyses of sand fraction (>0.7 mm) of samples from wells drilled by J. M. Zeigler and associates, supplemented by unpublished notes by them and archived at the Woods Hole Oceanographic Institution; well 1 is from LeBlanc et al. (1986, Sheet 1). See Figure 2 for locations of wells and Figure 7 for lithology legend.

(1965) as fluvio-glacial in origin. If the clay does represent such a depositional environment, it is probably an outlier rather than a westward prolongation of the Truro plain as inferred by Zeigler et al. (1965, their Fig. 2). Possibly the clay is not fluvioglacial but rather lagoonal in origin, deposited in the Provincetown Hook complex during the Holocene. The sediments between the Provincetown Hook facies and the lake deposits in borehole C at Stark’s display less wear than the Provincetown Hook marine sediments above, leading Zeigler et al. (1965) to propose that they are of fluvial-glacial rather than marine ori-

Late Quaternary construction of Cape Cod, Massachusetts

13

Figure 10. Lithology of wells drilled on the Provincetown Hook. Description of wells A, C, AA, and BB drilled by J. M. Zeigler and associates is based on the microscopic analyses of the sand fraction (>0.7 mm) of samples supplemented by unpublished notes by Zeigler and associates. Lithology of well B is based on notes transcribed by H. Hoskins of the Woods Hole Oceanographic Institution (WHOI) (included with the unpublished notes of J. M. Zeigler archived at WHOI) from an unpublished manuscript by M. F. Knout in the U.S. Geological Survey’s Boston office. See Figure 2 for locations of wells and Figure 7 for lithology legend.

gin. Zeigler et al. (1965) also proposed that this sequence at a depth of 40 m probably is in place and marks the most westerly occurrence of the Truro drift. It probably defines the edge of an erosional terrace cut in front of the easterly retreating High Head sea cliffs. Possibly the sediments also may represent sediment slumped as the cliff retreated eastward during the Holocene. On both the Stark’s (C) and the Provincetown Power Plant (B) boreholes is a massive black mineral, which may be wad. This mineral generally is found in marshy areas and is formed as a result of the decomposition of manganese minerals. In borehole B the mineral is found in sediments inferred to be lacustrine in origin, whereas at site C the wad is in association with reworked glacial sediments. At site B the lacustrine and the Provincetown facies are separated by an unfossiliferous sequence of hard clay and coarse sand. Part or all of this interval may be equivalent to the amorphous seismic layer described

by Oldale (1988) from Cape Cod Bay (Fig. 22). At Race Point Spit (borehole A) the Provincetown Hook facies rests on a pebbly, poorly sorted medium sand with water-worn clasts that also may be equivalent to the amorphous layer offshore. These sediments, like the fine sediments at Stark’s (borehole C), were previously assigned by Zeigler et al. (1965) as Eocene in age; we suggest that they are late Pleistocene deposits. Eastern offshore area. The narrow shelf east of the lower Cape is made up of a wave-built terrace south, and a wave-cut terrace north, of Chatham (Fig. 2). Whereas the wave-built terrace surface sediments consist of coarse sand, the wave-cut terrace is mantled by a lag deposit of reddish coarse sand and gravel (Schlee and Pratt, 1970; Schlee, 1973; Aubrey et al., 1982). Clasts recovered from the terrace’s seaward scarp during the current investigation were found to be coated with manganese oxide, suggesting that little or no sediment is being

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deposited on the scarp at present. High-resolution seismic reflection profiles taken in the wave-built terrace with an EG&G Uniboom display a prominent reflector about 10 to 30 m below the sea floor, an unconformity shoaling to the north and south (Fig. 11) (Aubrey et al., 1982). Cutting this surface is a channel system draining to the east. Below the unconformity is an acoustically massive unit and above it a layer displaying a complex textural fabric. The upper sediment unit in places lacks internal reflectors, in others it has an extensive network of discontinuous reflectors, and in still others it displays foreset bedding dipping to the south and west with a source to the northeast. The surface of the upper unit is channelized, resulting in an undulating topography. These channels, which

become more numerous and narrower toward shore and terminate at a depth of about 18 m, may have been produced by rip currents (Dillon et al., 1979). Aubrey et al. (1982) identified the subsurface horizon as the top of the depression occupied by a protrusion of the South Channel lobe at the time that the outwash sediments to the north and west of the Nauset Marsh were deposited. The channels on the surface supposedly were eroded subaerially, possibly from discharge from Cape Cod Bay Glacial Lake, before the deposition of Holocene marine sediments filled the low. Aubrey et al. (1982) also postulated that the sediments above the unconformity depression represent either outwash sediments deposited by the South Channel lobe or Holocene marine sediments derived from the eroding cliffs

Figure 11. Single-channel seismic reflection profile of the northern edge of the sediment-filled low seaward of Cape Cod’s elbow. The undulating prominent reflector (UR) pinching out on the platform’s south scarp was assumed by Aubrey et al. (1982) to be the base of the low left behind by a sublobe of the South Channel lobe when it retreated eastward. They determined that this low had a relief of about 0.03 s (25 m, assuming a sediment velocity of 1,600 m/s). Seismic reflection profiles taken with a more powerful seismic source (see Fig. 12) demonstrate that this reflector is the top of a thick section filling a much deeper low whose foundation is basement. This low has a relief exceeding 0.1 s (more than 80 m, assuming a sediment velocity of 1,600 m/s). The thick unit below the undulating reflector is believed to have been deposited by the retreating sublobe and the unit above the reflector by marine processes during the Holocene. Profile from the archives at the U.S. Geological Survey, Branch of Atlantic Marine Geology, Woods Hole, was recorded by D. Twichell and D. Aubrey during R/V Neecho cruise NE-06 in 1979 using an EG&G Uniboom. Profile 11 in Figure 1. M = acoustic multiple.

Late Quaternary construction of Cape Cod, Massachusetts to the north. If these cliffs eroded in the past at the present rate of 0.8 m a–1 (Zeigler et al., 1964a), they would have been capable of providing sufficient detritus to fill the low. We suggest, however, that the lower massive unit is of glacial origin and was deposited by the lobe of the South Channel lobe, and that the upper unit is a Holocene marine sequence made of sediments derived from the north. Seismic reflection profiles taken using a 40-in3 air gun as a sound source show that the massive lower unit is quite thick, filling a depression more than 100 m deep (Fig. 12). Its base is formed by basement, its northern rim by the edge of the erosional terrace to the north, and its southern one by a scarp defining the northern termination of the coastal plain sediments. Prior to the formation of the erosional terrace and the deposition of the upper unit, the northern side of the depression may have risen 60 m above the general level of the depression. Above the massive sediments filling most of the depression are the marsh deposits behind the Nauset Beach on the west, the complex bedded unit on the shelf in the center, and Fan B on the outer shelf on the east, a depocenter pinching out on the glaciomarine sediments of the Wilkinson Basin complex in the western Gulf of Maine (Fig. 13; see also 25E). The massive lower unit represents sediments deposited by the retreating sublobe and sediments that collapsed into the low from the retreating ice, and represents a southeastern extension of the Eastham plain deposits. The discontinuous patches immediately above basement may be part of this fill or may represent coastal plain erosional remnants. The rest of the low’s fill, the complex bedded unit, and Fan B (Fig. 13) consist of sediment eroded from the glacial terrains to the west and from the wavecut terrace north of the depression. Although a considerable volume of glacial and marine sediment has accumulated in it, the depression still has a topographic expression as evidenced by the broad reentrant on the slope’s contours (Fig. 1). Once the low was nearly filled, nearshore processes formed the Nauset and Monomoy bars, and the filled low was cut in two. The sediment regime seaward of the barrier island has remained a highenergy environment as documented by the rip current channels and large waves, whereas the segment of the low landward of the barrier became a low-energy domain, allowing the development of the present marshes. The wave-cut terrace fronting the sea cliffs of the lower Cape is 2.8 to 7.0 km wide and 33 km long with its outer edge at a depth of 38 to 40 m (Fig. 2). Parts of the inner boundary of the terrace are buried by onlapping sediments and along its southern rim is at least one sediment-filled channel that probably drained southward to the low south of the terrace. The terrace’s surface is rough with subtle changes in its seaward gradient probably documenting temporary sea-level stands or changes in rates of sea-level rise during its erosion. The slope along its seaward edge has a relief of 81 to 99 m and a gradient of about 5° (assuming a sediment velocity of 1,600 m s–1; Fig. 14). This scarp defines the westerly boundary of the South Channel lobe at the time that the outwash plains of the lower Cape were

15

Figure 12. Line interpretation of a single-channel seismic reflection profile of the sediment-filled low seaward of Nauset Marsh. The profile trends obliquely across the filled low, causing the apparent northerly dip of the sea floor. The low is bordered by coastal plain sediments to the south and outwash sediments to the north. Prior to the construction by erosion of the terrace to the north, this side of the low may have had a relief of about 90 m. Erosional remnants within the low may consist of either coastal plain or glacial drift sediments. Much of the low’s fill was deposited by an easterly retreating sublobe of the South Channel lobe. The rest of the fill, about 20 to 30 m thick, was derived from the north when a terrace was eroded during the latest Wisconsin–Holocene transgression from lower Cape outwash sediments. Line interpretation is of a profile from the archives at the Woods Hole Oceanographic Institution recorded by R. N. Oldale, E. Uchupi, and K. E. Prada during R/V Gosnold cruise 146 in 1969. Sound source: 40 in3 air gun. Profile 12 in Figure 1.

Figure 13. Single-channel seismic reflection profile of the seaward edge (Fan B) of the sediment-filled low seaward of Cape Cod’s elbow. B = basement; G = glacial sediments; H = Holocene marine sediments. Note seaward-dipping reflectors at the northwest end of the seismic section. Profile from the archives at the Woods Hole Oceanographic Institution was recorded by E. Uchupi and R. D. Ballard during R/V Gosnold cruise 204 in 1973. Sound source: 40 in3 air gun with bubble suppressor. Profile 13 in Figure 1.

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Figure 14. Single-channel seismic reflection profile of the eastern edge of the erosional platform north of sediment-filled trough in Figure 11. B = basement; D = debris flows derived from the slope to the west; E = escarpment defining the western edge of the sediment wedge deposited by the South Channel lobe that created the lower Cape; G = glacial sediments; H = Holocene deposits; M = multiple reflection, P = platform. The platform was formed by wave erosion of the proximal end of the outwash sediments during the latest Wisconsin–Holocene transgression. Profile is from the archives at the Woods Hole Oceanographic Institution and was recorded by E. Uchupi and R. D. Ballard during R/V Gosnold cruise 204 in 1973. Sound source: 40 in3 air gun with bubble suppressor. Profile 14 in Figure 1.

deposited. Sediments forming the core of the terrace are acoustically homogeneous and are about 0.12 s (96 m, assuming a velocity of 1,600 m s–1) thick. Sediments beyond the outer scarp consist of an irregular surface lower unit with a maximum thickness of 60 m resting on basement and terminating on a bulge. This unit may be a debris flow that originated in the scarp to the west when the South Channel lobe retreated away from the scarp, leading to its partial collapse. Above the debris flow sequence is 30-m-thick Fan C consisting of detritus from the erosional terrace. At its distal end the fan onlaps the Wisconsinan glacial marine sediments of the western Gulf of Maine. Cape Cod Bay. Basement in Cape Code Bay, exposed along its northwest part, slopes gently toward the northeast (Oldale et al., 1973; Oldale and O’Hara, 1990). Two drainage systems are carved on this surface, one northwest draining the area via a water gap north of Provincetown, and a southern one draining via a channel in the outer lower Cape along an east-west–trending Late Triassic–Early Jurassic rift (Oldale and Tuttle, 1964; Ballard and Uchupi, 1975). Above the basement are isolated erosional remnants (Fig. 18) that have been inferred as coastal plain sediments comparable to those described from Marshfield by Kaye (1983). Resting on the older units are lacustrine sediments (originally interpreted as Cenozoic sediments by Hoskins and Knott, 1961) deposited in Cape Cod Bay Glacial Lake during the retreat of the Cape Cod glacial lobe. The lacustrine facies along the margins of the bay consist of outwash deltas. In the center of the bay are well-stratified ice-contact sediments, which Oldale (1988) believed were deposited in a series of fans by meltwater that entered Cape Cod Bay Glacial Lake along the

base of the retreating ice. On the northern end of Cape Cod Bay the lacustrine sediments are separated from a well-stratified acoustic unit by a prominent reflector documenting a rapid and drastic change in sedimentation. Seismic profiles extending from two coring sites in Stellwagen Basin (Tucholke and Hollister, 1973) to Cape Cod Bay indicate that the sediment above the horizon is glaciomarine. Oldale (1988) and Oldale et al. (1993) have proposed that the lower glaciomarine sediments were deposited when the ice retreated north of Cape Cod Bay, as far as northeastern Massachusetts prior to 14,000 B.P., and the upper sediments 14,000 to 13,000 B.P. when the ice retreated north of Boston and relative sea level near Boston was +33 m higher than now. Above the glaciomarine and lacustrine sediments in northern Cape Cod Bay is a wedge-shaped acoustic amorphous layer thickening to the south, which vibracore sampling indicates consists of sand (Fig. 22). Oldale (1988) suggested two origins for the layer. In one the amorphous unit was deposited in a marine environment, an origin supported by the foreset bed structure displayed by the unit west of Provincetown. This delta supposedly was created when the Cape Cod Bay region rebounded isostatically due to ice unloading, relative sea level dropped to –43 m below its present level, and Cape Cod Bay was eroded by fluvial processes. Sediments forming the delta were derived from this fluvial erosion of Cape Cod Bay. In the second origin the unit was emplaced by a submarine landslide or debris flow, an origin suggested by the amorphous acoustic signature of the layer. If the layer is a debris flow, it may have originated in the High Head sea cliff and may be more extensive than indicated by Oldale (1988). If the debris

Late Quaternary construction of Cape Cod, Massachusetts flow originated on the cliff near the terrace at Stark’s (borehole C: Fig. 10), the fine-to-coarse sand and compact clay on Provincetown at depths of 37 to 55 m below sea level above the lacustrine sediments (borehole B), and the pebbly, silty poorly sorted sand on Race Point at a depth of 60 m below sea level (borehole A) may be onshore equivalents of this acoustic unit. Birch (in Oldale and O’Hara, 1990) suggested a third possibility, proposing that the amorphous layer is an ancestral Provincetown Hook formed during the late Wisconsinan submergence of the outer lower Cape. The Wisconsinan sediments of Cape Cod Bay and the valley fill above the late Wisconsinan regression are truncated by a wave-eroded surface formed during the Holocene transgression that began about 12,000 B.P. Above this hiatus are Holocene marine sediments deposited during and subsequent to this transgression. Billingsgate Shoal. Billingsgate Shoal is a triangularshaped topographic high 14 km long by 6 km wide, that is asymmetric (steeper toward the southeast) and extends southwestward from Wellfleet Harbor. According to Davis (1896), the development of the present “west concave” shoreline of lower Cape Cod began in concert with the growth of the Provincetown Hook. Prior to that time the western (Truro) coast of the lower Cape had been shaped into a long, straight shoreline extending from the relic wave-cut cliffs of High Head southward to the tombolo now forming the western shore of Wellfleet Harbor. Davis was aware of suggestions that outlying islands may once have stood on the site of Billingsgate Shoal, and he wondered whether the present southwest-trending nearshore shoals at the base of Billingsgate Shoal represent a new feature produced by the post-Provincetown Hook excavation of the concave shoreline, or rather were somehow related to earlier land forms on Billingsgate Shoal. Giese (1963) developed the second alternative theory and in so doing suggested that the present tombolo-formed shore of Wellfleet Harbor is a recent feature, and that the pre–Provincetown Hook shoreline, far from being straight, as Davis had postulated, curved out toward the southwest following the general outline of the present Billingsgate Shoal. A comparison between U.S. Coast and Geodetic Survey surveys of 1849–1850 and 1934 indicated that the shoal is slowly migrating southeastward, and that more than 10 million m3 of sediment were removed from the shoal during the 85-yr period. From the trend of the wave-cut cliffs of High Head and the depositional history of the Provincetown Hook, and assuming the geodynamic processes proposed by Davis (1896), Giese suggested the following origin for Billingsgate Shoal. Prior to the formation of Provincetown Hook about 6,000 B.P., the western shore of the lower Cape was eroded by waves from the northwest, and the resultant strong southerly littoral drift deposited this debris to form a feature comparable to Sandy Neck along the bay’s south shore. As the Provincetown Hook was formed, the region was more protected from northerly waves, the southerly drift diminished, subaerial Billingsgate Shoal began to both erode and submerge in response to a rising

17

sea level, and eventually the spit diminished to its present form. Evidence supporting this model for the development of the shoal is the cordon of beaches connecting glacial sediment outliers landward of the shoal in Wellfleet Harbor; these beaches were not in existence in the 18th century. Their presence supports the contention that the littoral drift is a recent phenomenon initiated after Billingsgate Shoal to the west began to waste away. Comparison of surveys also indicates that the irregular southeast shore of the currently eroding shoal has become smoother with time, a feature to be expected as the shoal erodes. Giese (1963) further speculated that Billingsgate Shoal may have originated as a ridge of glacial deposits. Oldale and O’Hara (1984) suggested that Billingsgate Shoal originated as a moraine formed by a minor advance of the Cape Cod Bay lobe, which deformed the Cape Cod Bay Glacial Lake sediments in front of it. Larson (1982) proposed that this readvance is equivalent to the one that formed the Monks Hill moraine west of Cape Cod Bay by the Buzzards Bay lobe (Fig. 22). Seismic reflection profiles recorded by Oldale and O’Hara (1990) (Figs. 15, 16B) demonstrate a complex pattern of internal reflectors dipping steeply toward the northwest in Billingsgate Shoal. This fabric appears to be of tectonic rather than of depositional origin and is probably due to the deformation of the glaciolacustrine deposits by ice pressure from the northwest (Fig. 16B). As the ice advanced southeastward, the sediments in front of it were thrust in that direction creating the shoal’s imbricate structure with the layers dipping toward the northwest. This tectonic stacking accounts for the ridge’s southeast flank being much steeper than its northwest one. The profiles indicate that the southwesterly tip of the moraine was planed off during the Holocene transgression, and the erosional surface was later buried by sediments as thick as 8 m, some of which have dipping internal reflectors (Fig. 15). These relict sediments, probably deposited by southwesterly flowing littoral drift as had been suggested by Giese, were interpreted by Oldale and O’Hara (1990) as relict bars or beaches deposited at lower sea levels. The rest of the moraine does not appear to have been affected significantly during the Holocene transgression with modifications consisting mainly of smoothing topographic irregularities by filling the lows along the crest of the ridge. The ridge probably extended farther east than it does now. How far it extends beneath the Wellfleet deposits of the lower Cape and whether it extended to the South Channel lobe are yet to be resolved. There is no doubt that before it was buried by Wellfleet plain deposits, the ridge must have influenced deposition not only from South Channel lobe east of Glacial Lake Cape Cod, but also from Cape Cod Bay lobe north of the lake. How long this influence persisted is another question awaiting answer. Although the ridge may have affected sedimentation in the region, it does not appear to have been a significant sediment source during its existence, as deposits do not appear to increase appreciably in thickness in the direction of the high. Stellwagen Basin. Sediments above basement on Stellwa-

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Late Quaternary construction of Cape Cod, Massachusetts Figure 15. Seismic reflection profiles of Billingsgate Shoal recorded by Oldale and O’Hara (1990) during their investigation of Cape Cod Bay. In contrast to the crest of the shoal, which only underwent some minor modification during the Holocene transgression (profiles 16 through 21), its southwest tip was extensively eroded (profiles 9 and 8) during the transgression. This erosional surface was later buried by sediment deposited by southwest-flowing littoral drift. See Figure 16 for locations of profiles. Profiles are from the archives at the U.S. Geological Survey, Marine Geology Branch, Woods Hole. Sound source: EG&G Uniboom.

gen Basin and east of Stellwagen Bank consist of coastal plain erosional remnants, glacial drift, Wisconsinan glaciomarine mud containing the foraminifera Elphidium clavatum (also excavatum) and Holocene marine mud (Oldale, 1988, 1989; Oldale et al., 1973; Schnitker, 1976). 14C age determinations from two cores raised from Stellwagen Basin yielded ages of 21,950 ± 1,350 and 18,900 ± 600 B.P. from the drift at the base of the cores at 20 and 22 m, and 13,130 ± 250 B.P. from the upper glaciomarine sediments at 13.5 to 13.8 m (Tucholke and Hollister, 1973). The dates of 14,250 ± 250 and 13,800 ± 300 yr reported by Kaye and Barghoorn (1964) from the glaciomarine mud in Lynn, Massachusetts, are now considered to be too young (R. N. Oldale, personal communication). Stellwagen Bank, a triangular-shaped topographic high located 10 km north of the tip of Cape Cod, may have had a

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geologic history similar to the lower Cape (Oldale, 1993; see also Ward, 1995, p. 15–19). The surface of the 40-km-long by 16- (southern end) to 14-km- (northern end) wide feature dips eastward. It, too, formed in an embayment between the Cape Cod Bay and South Channel glacial lobes, with the former being located somewhat west of Stellwagen Basin and the latter in contact with the east flank of Stellwagen Bank. Its easterly dip and westerly curve suggest that the bank formed from a westerly source, the Cape Cod Bay lobe. If so, it must have been emplaced prior to the deposition of the late Wisconsinan glaciomarine sediments in Stellwagen Basin. Oldale (1993), however, believed that the bank was deposited from the east by the South Channel lobe as documented by the coarsening of the bank’s surface sediment in that direction. He ascribes the bank’s present outline to postglacial changes that occurred during and subsequent to its drowning. The possibility that the bank may have been formed from both easterly and westerly sources and deposited in a marine bay between South Channel and Cape Cod Bay lobes also merits consideration. In contrast to the lower Cape, Stellwagen Bank was formed by deltas graded to a marine embayment rather than a freshwater lake. The bank also differs from the lower Cape in that much of the outwash deposits may rest mainly on coastal plain erosional remnants rather than basement (Oldale et al., 1973). Whereas a large segment of Cape Cod remained above sea level during the Holocene transgression, Stellwagen Bank was drowned and its

Figure 16. A, Mode of formation of Billingsgate Shoal by a readvancing glacial lobe. Densely stippled pattern on the flanks of the shoal represents lacustrine sediments deposited after formation of the shoal. Adapted from Oldale and O’Hara (1984, Fig. 12). B, Topographic map of Billingsgate Shoal showing locations of seismic reflection profiles in Figure 15. Distribution of beach and bar deposits at the southwestern end of the shoal is from Oldale and O’Hara (1990, Fig. 10). Contours in beach and bar deposits are isopachs in meters.

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top truncated. Much of the resulting detritus was transported westward, possibly forming the high parallel to its western flank reversing its dip; some of the detritus spilled from the bank into Stellwagen Basin beyond the high. Western Gulf of Maine. Basement in the Gulf of Maine consists of Paleozoic and older rocks on which is faulted a Mesozoic rift system (Ballard and Uchupi, 1975). Above the basement are a few isolated coastal plain erosional remnants such as Fippenines Ledge at 69°18′W, 42°48′N with its Eocene siliceous sediments (Schlee and Cheetham, 1967). Mantling the floor of the Gulf is a thin veneer of Pleistocene sediments that document the late Wisconsinan glaciation and subsequent glacial retreat. In most of the basins in the Gulf of Maine the glacial sediments are covered by an acoustically transparent mud deposited during the Holocene (Schlee, 1973; McClennen, 1989). Except for a few scour depressions near topographic highs and possible pockmarks associated with biogenic gas seeps, these Holocene muds display a smooth nearly level surface. At the base of the glacial sequence is a poorly sorted drift (Scotian Shelf Drift) (King and MacLean, 1976; King, 1980) containing an appreciable amount of gravel, a unit that is particularly thick in Great South Channel (Knott and Hoskins, 1968; Uchupi, 1970). The basal till in Great South Channel contains reworked Paleocene and Eocene bryozoans, Pleistocene foraminifera, glauconite, shell fragments and fish teeth. 14C dating of total carbon from the deposit by Bothner and Spiker (1980) provided an age no older than 20,000 B.P. for the till. Reworking of the till during the late Wisconsin-Holocene transgression led to the formation of wide expanses of lag gravel in the Gulf of Maine (Schlee, 1973; Schlee and Pratt, 1970). Above the drift is the Emerald Silt, a glaciomarine fine-grained, wellbedded locally gravely silt sand deposited in brackish-to-normal marine conditions (Belknap et al., 1991). Schnitker (1988) stated that the marine incursion into the Wilkinson Basin in the western Gulf of Maine had started by at least 17,600 B.P. and that glaciomarine conditions ended 14,000 B.P. in the outer Gulf and 12,000 B.P. near the present coast. Two models have been proposed for the deposition of the glaciomarine sediments in the Gulf of Maine, an ice shelf (Hughes et al., 1985; King and Fader, 1986; Dyke and Prest, 1987; Belknap and Shipp, 1988; Schnitker, 1988) and a grounded subpolar marine-based glacier (Oldale et al., 1990). The wide distribution of seismic layers in the glaciomarine silt, the presence of clasts in the silt, and the absence of a microflora are believed by Schnitker (1983, 1988) to be compatible with deposition beneath an ice shelf. Previously Schnitker (1986) suggested that the ice shelf occupied all the Gulf of Maine being grounded near the coast with its outer edge resting on Georges Bank. More recently, Belknap et al. (1991) proposed that the ice lifted off the basins in the Gulf of Maine in sequential order from south to north with “till-tongues” (wedges of acoustical reflections interbedded with the glaciomarine sediments; Mosher et al., 1989; Piper et al., 1987) along the basin margins documenting ice-grounding effects and later furrowing

by icebergs. Oldale et al. (1990) rejected the ice-shelf model, pointing out that there are no modern examples of a temperate shelf, that temperate conditions that existed at that time in southern New England would have led to the destruction rather than the preservation of an ice shelf, and that presence of relatively warm ocean water (5° to 8°C) beneath the ice during late Wisconsin (Schnitker, 1976) also is incompatible with an ice shelf. They also pointed out: (1) if the mud was deposited beneath an ice shelf, then its repetitive rhythm deposition was created by repetitive and simultaneous melt-out of the ice shelf, an unlikely event; (2) basal melt-outs on modern ice shelves occur only near the grounding line shortly after flotation and that the ice shelf seaward of the grounding line is free of basal debris; and (3) such melt-outs produce only a diamicton, not a glaciomarine silt. For these reasons, Oldale et al. (1990) proposed a grounded marine-based glacier model. Glaciomarine deposition in this model is from sediment-laden meltwater streams that reach the open sea through tunnels at the base of the gounded ice front in the manner described by Pfirman (1985) for the Svalbard glacier. These streams deposit their coarse load near the grounding line, producing features similar to “ice tongues.” The finer sediments were transported farther seaward by near-sea-surface plumes that were active along the ice front. Deposition from these plumes, which are capable of transporting muds as far as 50 km offshore (Molnia, 1983), is responsible for the widespread rhythmic stratification of the glaciomarine sediments. GLACIAL CHRONOLOGY AND SEA LEVEL The sea-level curves described by Fairbanks (1989) and Bard et al. (1990) show sea level rising at about 18,000 B.P. and display no evidence of when it bottomed out. Thus, the maximum southern extent of the Laurentide glaciation must have occurred earlier. Similarly, radiogenic dating (35,000 to 20,000 B.P.) of a till in Great South Channel suggests that the Laurentide ice reached its maximum extension in New England about 20,000 B.P. (Bothner and Spiker, 1980). More recently Oldale (1989) proposed that full glaciated conditions in southern New England probably took place between about 21,000 and 18,000 B.P., with retreat from the glacial maximum beginning 18,000 B.P. As the ice advanced southward it picked up and incorporated into the drift organic remains with the youngest material dated, about 21,000 B.P., representing the time of the advance of the Laurentide ice sheet to its southern terminus at the mouths of Northeast and Great South channels, the northern edge of Georges Bank, and the terminal moraines in Martha’s Vineyard and Nantucket Island. It is because of the incorporation of old material into the drift that the dates of 20,000 and 26,000 B.P. on shell and wood from the Wellfleet and Highland plain deposits (Zeigler et al., 1964a; Oldale, 1968) and other similar dates are an indication only that the drifts in Cape Cod are not older than late Wisconsin, rather than an indication of their absolute age.

Late Quaternary construction of Cape Cod, Massachusetts An age of 21,000 B.P. for the time when the Laurentide ice sheet reached its southern terminus is also suggested by the age of one of the Heinrich layers (rich in ice-rafted debris and poor in foraminifera) in the North Atlantic. These layers were deposited as a result of the rapid advance of ice streams in eastern Canada and maximum calving along the ice front as it reached its maximum seaward position (Bond et al., 1992). A date of 20,000 B.P. for maximum Laurentide glaciation also is suggested by ages from other sites along the southern margin of the Laurentide glacier on land. Sediments indicative of a cool climate from western Long Island and deposited prior to the late Wisconsin glaciation, for example, yielded ages ranging from 27,950 ± 450 to 21,750 ± 450 B.P. (Sirkin and Stuckenrath, 1980). According to Stone and Borns (1986), the 21,750yr date is the youngest subtill date from the New England– Long Island region. A stump cluster in growth position near Cincinnati that was killed by the overriding Laurentide glacier yielded an average age of 19,670 ± 68 B.P. (Lowell et al., 1990). The maximum extent of the Lake Michigan lobe has been dated as 19,500 ± 200 and 20,000 ± 200 B.P. and the Fayette maximum as 21,000 B.P. (Broecker and Denton, 1990, and references therein). In the reconstruction of the paleogeography of the Cape Cod region discussed below, we assume that the Laurentide ice sheet reached its maximum southerly position 20,000 B.P., and that retreat began soon after, to account for the marine conditions that were established in the western Gulf of Maine no later than 17,600 B.P. Glacial retreat in southeastern Massachusetts was relatively rapid as marine conditions were established in Wilkinson Basin in the western Gulf of Maine no later than 17,600 B.P. (Schnitker, 1988) and Stellwagen Basin northeast of Boston no later than 14,000 B.P. (Tucholke and Hollister, 1973; Oldale et al., 1993). By 14,000 B.P. the ice front in New England had retreated from its maximum southern position to a point some 200 km to the north, a retreat rate of 50 km/1,000 yr. If the ice retreated at a uniform rate (a questionable assumption), then Cape Cod was ice free by no later than 15,000 B.P. Such a scenario is supported by studies in coastal Maine by Smith (1982, 1985), which indicate that deglaciation of the Gulf of Maine took place between 17,000 and 13,000 B.P. This rapid glacial decay in late Wisconsin was due to an increase in summer insulation in the Northern Hemisphere (Lehman et al., 1991) which, according to Fairbanks (1989), was underway by about 18,000 B.P. and peaked at a radiocarbon date of 12,000 B.P. Deglaciation appears to have been in two major steps separated by a period of little or no ice volume loss. Rates of global glacial meltwater discharge calculated from the Barbados sea-level curve by Fairbanks (1989) indicated that periods of maximum melting took place at about 12,000 and 9,500 B.P.; and the period of minimal glacial meltwater discharge, the Younger Dryas event, occurred between 11,000 and 10,500 B.P. Associated with the latest melting event was the invasion of warm, salty Indian Ocean water into the North Atlantic via the Agulhas-Benguela Current system about 9,000 B.P. According to

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Jones (1996), the invasion of this warm, salty water led to stabilization of the present climatic and oceanic regime of the North Atlantic. Major meltwater discharge into the Gulf of Mexico was more sporadic, having occurred from 16,500 to 15,500 B.P. and from 15,250 to 14,900 B.P., with minor discharges claimed to occur at 14,400, 14,100, 13,800, 12,800, 12,100, and 11,900 B.P. (Leventer et al., 1982). The pronounced decrease in discharge into the Gulf of Mexico between 11,000 and 10,000 B.P. was the result of a change in discharge direction from southward via the Mississippi River to eastward via the Hudson and St. Lawrence Rivers (Broecker et al., 1988; Joyce et al., 1993). During this rather abrupt retreat disrupted by periodic readvances (such as the Cary readvance between 15,000 and 14,000 B.P.), when the ice deformed the sediments in front of it, the glacial morphology of the Cape Cod region was created. The morphologic development of the glacial Cape may have been influenced by the Younger Dryas event when reoccurrence of glacial conditions in the western Gulf of Maine (Wilkinson Basin) disrupted the nonglacial environment of the region. This refrigeration was the result of a large and sudden discharge through the St. Lawrence River with the meltwater reaching the Gulf of Maine via the Nova Scotian Shelf (Schnitker et al., 1991). However, no evidence for this renewed glacial activity has been observed on the Cape. Another factor that influenced the development of Cape Cod’s morphology was the latest Wisconsin-Holocene rise in sea level. Although it is difficult, if not impossible, to construct a global eustatic sea-level curve for the last deglaciation because of local neotectonism (see Emery and Aubrey, 1991), the calibration of the 14C time scale curve from Barbados (Fairbanks, 1989; Bard et al., 1990) (Fig. 17) is probably a close approximation of the response of sea level to Wisconsin glacial decay. At the glacial maxima about 18,000 B.P. , sea level was about 130 m below its present level. From there sea level rose at a rate of about 11 m/1,000 yr, reaching a depth of 20 m below the present level 8,000 B.P. Approximately 6,000 B.P., global mean sea level was at a depth of 10 m and 5,000 B.P. at about 7 m below its present level. Glacioisostatic adjustment in the Cape Cod region complicates local relative sea-level history, so global isostatic curves are of limited use here. The curve compiled by Milliman and Emery (1968) for the nonglaciated shelf off the eastern United States tends to follow the general trend of the curve of Bard et al. (1990), but those constructed for New England depart noticeably from the Barbados trend. As a result of crustal depression by glacial loading, relative sea level in northeastern Massachusetts was +33 m higher than its present level about 14,000 B.P. This high-level stand, documented by emerged glaciomarine sediments, raised paleoshorelines and ice-contact deltas (Oldale, 1985b; Oldale et al., 1993, and additional references therein) ranged from +18 m in Boston Basin to +130 m in central Maine. The age of the highstand ranges from 14,250 ± 250 and 13,800 ± 300 B.P. in Lynn, Massachusetts (Kaye and Barghoorn, 1964), about 14,250 B.P. in the Merrimack River valley, Massachusetts (Oldale et al.,

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Figure 17. Sea-level curve during the last deglaciation from Barbados, Caribbean, and the Cape Cod region. The Barbados curve is from Bard et al. (1990, Fig. 1). The curves from Cape Cod are from Redfield and Rubin (1962, Fig. 2), Barnstable; Kaye and Barghoorn (1964, Fig. 5), Boston; Oldale and O’Hara (1980, Fig. 2), southeastern Massachusetts shelf; and Oldale et al. (1983, 1987, 1993, Fig. 5), northeastern Massachusetts and the western Gulf of Maine. Date from TT site in Nantucket Shoals (see Fig. 1 for location of site) is from Groot and Groot (1964). The curves of Kaye and Barghoorn and Oldale et al. have been modified to indicate the possible maximum drop in sea level during the 12,000 B.P. regression due to crustal rebound.

1993), and 13,830 ± 100 B.P. in coastal Maine (the oldest date from the glaciomarine Presumpscot Formation) (Smith, 1985). Ages of 13,670 ± 145 and 14,090 ± B.P. for the glaciomarine muds off New Hampshire and Maine (Birch, 1990; Kelley et al., 1992) lend additional support to an age of about 14,000 B.P. for the transgression. Cape Cod, Nantucket Sound, and the southern end of Cape Cod Bay were not inundated during this marine transgression as isostatic rise in the peripheral bulge due to glacial unloading in adjacent regions exceeded the rate of eustatic rise in sea level prior to 16,000 B.P. This high stand was short-lived, as the crust rebounded rapidly when its glacial ice load was removed. As the crust rose, sea level dropped and by 12,000 B.P. it was at a level –43 m below its present one off northeastern Massachusetts (Oldale et

al., 1993). Depth estimates during this low stand range from –22 m in Boston (Kaye and Barghoorn, 1964), –31 m off New Hampshire (Birch, 1990), –55 to –60 m off Maine (Belknap et al., 1989; Shipp et al., 1989), and –40 m off Penobscot Bay, Maine (Knebel and Scanlon, 1985). The sea-level elevation estimated by Kaye and Barghoorn (1964) for Boston, which is based on dates from freshwater peats at elevations ranging from –8.3 to –0.2 m below present sea level, is probably too shallow; an elevation of –30 m below present level may be more realistic (Fig. 17). Similarly, the –43-m elevation estimated by Oldale et al. (1993) by drawing a smooth curve through the points also may be too shallow; an elevation of about –55 m below present level is obtained if the curve is drawn through the Merrimack paleodelta and the Jeffreys Ledge barrier spit points (Fig. 17). Oldale et al. (1993) also proposed that the depth of –31 m for the New Hampshire shelf (Birch, 1990) may be too shallow pointing out that the truncated tops of the drumlins east of the Isles of Shoals (Birch, 1984) indicate a low stand of –55 to –60 m. By redrawing the sea-level curve of Birch (1990), an elevation of about –40 m is obtained. Even with these corrections these depths are shallower than the values displayed by the Barbados 14C sea-level curve where depths range from –72 m 12,000 B.P. and –63 m 11,000 B.P. below the present level. If this difference is real, then from 12,000 to 11,000 B.P. the crust in New England was still not in equilibrium but was rebounding, and it was only much later, about 6,000 B.P., that the crust in the region became relatively stable. A slowing of crustal uplift coupled with an increase in eustatic sea-level rise 12,000 B.P. caused the sea level to rise again in New England (Oldale et al., 1993). Such an increase in meltwater discharge 12,000 B.P. is supported by the Barbados sea-level curve (Fairbanks, 1989). The postglacial low stand in New England not only becomes deeper toward the north (–43 m off Massachusetts and –65 m off Maine) but also younger. It is 12,000 B.P. off northeastern Massachusetts, 12,080 to 10,820 B.P. off New Hampshire (Birch, 1990; recalculated by Oldale et al., 1993, as 11,400 B.P.), and approximately 11,000 B.P. off Maine (Kelley et al., 1992). Features associated with this sea-level regression include a paleodelta off the Merrimack River (Oldale et al., 1983), a submerged barrier beach and lagoon on Jeffreys Ledge (Oldale, 1985a), the seaward limit of shelf valleys off New Hampshire (Birch, 1984), a paleodelta off the Kennebunk River in Maine (Belknap et al., 1989), submerged terraces off Maine (Shipp et al., 1989), and a regressive unconformity in Penobscot Bay, Maine (Knebel and Scanlon, 1985). The magnitude of this crustal rebound and the intense deformation of the coastal plain and drift sediments by ice pressure in Gay Head, Martha’s Vineyard, are clear indications that the Laurentide ice sheet was appreciably thick at its southern terminus in New England. The peripheral bulge associated with this ice was located an unknown distance from its margin, perhaps on the mid- to outer shelf south of Martha’s Vineyard and Nantucket Island and mid- to outer shelf of Georges Bank. Dillon and Oldale (1978) proposed that subsidence of this bulge caused the

Late Quaternary construction of Cape Cod, Massachusetts shoreline features formed during the Holocene transgression on the mid- to outer shelf to be downwarped northeast of an inflection zone off central New Jersey. However, the lack of inundation of Nantucket Sound during the 14,000 B.P. transgression suggests migration of the peripheral bulge through time, as expected by physical arguments. In Boston, sea level has been reported to be within –0.6 m of its present level by 2,800 B.P. and about –0.5 m by 1,700 B.P. (Kaye and Barghoorn, 1964). However, Redfield (1967), using cores from the Neponset River, found sea level to be about –6 m at 2,800 B.P., and –3 m at 1,700 B.P. Redfield (1967) discounted the Kaye and Barghoorn (1964) measurements since they were taken in filled marshland, where compaction likely affected the vertical position of the samples. In Barnstable, along the north shore of Cape Cod Bay, sea level reached an elevation of –5 m at about 3,000 B.P. and –2 m at 2,000 B.P. (Redfield and Rubin, 1962). Tide gauge data indicate that coastal New England is currently subsiding relative to sea level at rates ranging from about 2 mm a–1 in Massachusetts to approximately 3 mm a–1 in Maine (Emery and Aubrey, 1991, p. 134; Gehrels and Belknap, 1993). Thus, subsidence today increases toward the direction of greatest isostatic crustal collapse of the peripheral bulge due to glacial unloading. This land subsidence relative to sea level is due to a combination of eustatic rise in sea level and tectonic processes, a neotectonism that may be the result of migration of the peripheral bulge in the direction of the retreating ice front. As the Laurentide ice sheet expanded southward, its weight caused the crust to subside and subcrustal flow to take place in the direction of glacial growth. This flow led to the formation of a peripheral bulge, whose relief reflected not just the weight at the southern periphery of the ice sheet, but the total ice load on the crust. As the ice sheet retreated northward, the bulge migrated with it and the formerly depressed crust was uplifted. Because the relief of the peripheral bulge reflected the load of the whole ice sheet, the crust rose above isostatic equilibrium. As the bulge migrated farther northward, it became more subdued and the uplifted crust slowly subsided. That relative rise in sea level is slower in Massachusetts than in Maine indicates the crust in the former is closer to equilibrium, a consistent finding because Massachusetts was deglaciated before Maine. This subsidence is responsible for sea level being 1.8, 3.0, 2.7, and about 1 m higher than eustatic sea level in Boston 7,000, 6,000, 5,000, and 4,000 B.P. as reported by Kaye and Barghoorn (1964) (Fig. 5). Since 3,000 B.P. the subsidence rate has been slow, no more than 1 to 2 m ka–1, and cannot be detected in curves constructed from radiocarbon dates. This subsidence, together with an eustatic rise in sea level of 1 to 2 mm a–1, is responsible for the present relative rise in sea level in the region of 2 to 3 mm a–1. A similar trend also was noted by Clark (1980) in his numerical models of sea-level changes on a viscoelastic Earth. According to Clark, relative change in sea level since the Wisconsin was not constant because of the effect of ice/water loading on the surface of the Earth and the geoid. He used a

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numerical model that included the above effects to predict relative sea-level change on a spherical viscoelastic Earth with realistic ocean and ice configurations and a realistic melting history. Using the model, Clark divided the Earth into five zones of relative sea-level changes. Zones I and II are pertinent to the above discussion. In Zone I, the glaciated area, the elastic and viscous uplift of the ocean floor, combined with the fall of the geoid, led to rapid uplift of the land relative to sea level. Zone II, the zone of the peripheral bulge, is characterized by submergence as a result of the collapse of the bulge. The transition between zones I and II is the region to which the peripheral bulge migrates, collapsing at the same time as it migrates. Clark (1980) found this transitional zone to be characterized first by emergence (uplift) and later by submergence (subsidence). Predicted relative sea-level curves for the Cumberland Peninsula, Baffin Bay, Canada, for example, indicate that emergence occurred soon after ice-sheet melting, and that this was followed by submergence resulting from the collapse of a migrating peripheral bulge. According to Clark, this submergence could not be due to a eustatic rise in sea level because he found no such rise during the past 5,000 yr. However, the sea-level curve from Barbados shows that sea level rose about 7 m during that time (Fairbanks, 1989, his Fig. 2). Relative sea level in southern New England today is rising at a rate of 2 to 3 mm a–1. Approximately 1 to 2 mm a–1 of this rise is believed to be due to a worldwide eustatic rise in sea level. If the remainder is due to continuing collapse of the peripheral bulge, then the effect of the collapse has persisted in southern New England for nearly 12,000 to 14,000 yr after the region was deglaciated. Such a persistence may not be realistic geologically. Relative rise of sea level in formerly glaciated regions, such as southern New England, may be due to other causes. Because it is not possible to isolate and quantify the causes of this rise in relative sea level and how far back in time these causes persisted, we use the trends displayed by the Barbados, Boston, and Barnstable sea-level curves in the discussions that follow. These curves indicate that sea level was at –40 m about 10,000 B.P., about –30 m 9,500 B.P., about –10 m 6,000 B.P., and that it approached its present position about 1,000 B.P. GLACIAL CONSTRUCTION OF CAPE COD The geologic investigations summarized above have yielded a coherent picture of the genetic geomorphic evolution of the Cape Cod region. There is no doubt that the Laurentide ice sheet and associated fluvio-glacio and glaciomarine processes played the major role in sculpting the Cape. Although this glaciation has overwhelmed features associated with older geologic events, the older features are still recognizable. Basement in the present site of the upper Cape is dominated by an east-west–trending high cut in two by a north-south valley extending from Cape Cod Bay to Nantucket Sound. In turn, the eastern basement segment is also nearly cut in two by a smaller valley. In the lower Cape, basement is scarred by an eastwest–trending Late Triassic–Early Jurassic rift (Fig. 18). Simi-

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lar rift basins also occur in the western Gulf of Maine. The topographic highs of the islands and associated lows north of the islands are a cuesta/lowland terrain carved out of the shelf strata by fluvial erosion during regressions that may date as far back as the Oligocene. Long and Block Islands, Martha’s Vine-

yard, Nantucket Island, and Georges Bank are segments of this west-east–trending cuesta, which extends from Long Island to the Grand Banks of Newfoundland via the outer banks on the Nova Scotian Shelf (inset map, Fig. 1). Landward of the cuesta’s north-facing scarp is a lowland floored by coastal plain

Figure 18. Pre-Wisconsin geologic map of the Cape Cod region. Compiled from Ballard and Uchupi (1975, Fig. 5), O’Hara and Oldale (1980, Fig. 4; 1987, Fig. 4), Oldale and O’Hara (1990, Fig. 5), and Oldale et al. (1973, Plate 2A). The erosional remnants identified in the map as coastal plain could be pre–late Wisconsinan glacial deposits. B.H. = Basement High.

Late Quaternary construction of Cape Cod, Massachusetts sediments immediately off the scarp and basement farther north. In Cape Cod Bay, north of the bay and east of the lower Cape are erosional remnants resting on basement that may be of coastal plain origin, although the possibility that they represent pre–late Wisconsin glacial deposits merits consideration. The water gaps of the fluvial system responsible for the formation of this terrain are represented today by Vineyard Sound, the valleys between Martha’s Vineyard and Nantucket Island, the low between the eastern end of Nantucket Sound and the lower Cape, Little Stellwagen Basin between the lower Cape and Stellwagen Bank, and Great South and Northeast channels on either side of Georges Bank. The fluvially carved terrain was probably modified by glacial processes during the Illinoian (Stage 6), and possibly even earlier glaciations, but the magnitude of these changes is yet to be resolved. After the Illinoian glaciation the region was inundated by marine waters as recorded by the deposits in Sankaty Head, Nantucket Island, and the marine sediments incorporated into the Wisconsinan glacial deposits of Cape Cod. The most significant change of the Cape’s fluvial terrain occurred during the Laurentide glaciation in the late Wisconsinan. The Laurentide ice sheet reached its maximum southern position about 20,000 B.P., a southern terminus defined by the southern limit of abundant gravel mapped by Schlee and Pratt (1970). The lobate ice front was composed of the Narragansett–Buzzards Bay lobe in the west, the Cape Cod Bay lobe in the center, and the South Channel, the largest lobe, in the east (Fig. 19). The viscoelastic models described by Peltier (1982) and Birchfield and Grumbine (1985) indicate that a trough with a relief of 200 to 300 m and a peripheral bulge with a relief of about 70 m were present beyond the ice front. This proglacial terrain morphology probably varied with time and space as a result of the variability of the ice sheet thickness in space and time, increase in glacier lobation, rapid fluctuations in the position of the ice front, and shifts in the centers of outflow (Teller, 1987). The peripheral bulge off southern New England was probably located at midshelf and possibly even farther south. An amphitheater-shaped depression on the outer shelf, the Mud Patch, and the gravitational structures on the continental slope may be the result of this crustal uplift. A lake may have occupied the 200- to 300-m-deep depression between the peripheral bulge and the glacial front in the area of the islands. Retreat of the lobes varied; the deposits associated with the western Narragansett–Buzzards Bay lobe are the oldest, and those associated with the eastern South Channel lobe are the youngest. Oldale (1992, p. 39) stated that the west-east differential retreat of the lobes may be due to the size of the basin occupied by the lobes, with the low occupied by the South Channel lobe being the largest and deepest. The subglacial channels in Nantucket Sound probably were carved out of the coastal plain strata at the time when the ice sheet front was at its maximum southerly position (Fig. 19). As the ice lobes retreated northward, the drift deposits of the islands and Nantucket Shoals began to accumulate with the drift in the

25

islands being graded to a possible lake that may have been present between the peripheral bulge and the ice sheet front. A readvance during the retreat deformed the coastal plain strata and outwash deposits to form the moraine in Gay Head, Martha’s Vineyard. Sometime prior to 18,000 B.P., the Buzzards Bay lobe retreated to a position west of the Elizabeth Islands, the Cape Cod Bay lobe to the northern edge of the upper Cape, and the South Channel lobe to just north of Great South Channel. Stages in this retreat are documented by the ice contact deposits along the south shore, the subsurface, and the north shore of the upper Cape. As the Cape Cod Bay lobe retreated northward, a lake began to form in the present position of Nantucket Sound, a lake that during its maximum development extended well into the upper Cape (see Profiles A-A′′ and B-B′: Fig. 6). This lake may be contemporary to the lakes that were formed in Long Island and Block Island Sounds during late Pleistocene glacial retreat (Lewis and Needell, 1987; Needell and Lewis, 1984; Needell et al., 1987). The lake in Nantucket Sound was dammed to the north by outwash deposits in the upper Cape, to the northeast by the ice-sheet front, to the south by the outwash deposits of the islands, and to the east by the outwash deposits of Nantucket Shoals (Fig. 20). Seeps from the proglacial lake eroded the valleys in the drift of Martha’s Vineyard and Nantucket Island, damming the lake to the south. The outwash plains in the upper Cape that were graded to the lake were built by a braided meltwater stream system draining southward from the Cape Cod Bay lobe. The outwash plain in the region of Nantucket Shoals was built from meltwater deposits originating in the South Channel lobe; these deposits pinch out atop the then subaerial continental shelf south of Nantucket Island. As the streams shifted their courses across the surface of the outwash plains, they built a complex of subdeltas superposed one atop another. As the sediments prograded southward over the lacustrine deposits, the subglacial valleys in Nantucket Sound were filled, subduing its topography. With continued sedimentation the lake became progressively smaller, but before the lake could be filled, it drained—probably to the southwest via Vineyard Sound and south via Muskeget Channel west of Nantucket Island. Prior to the formation of the Harwich plain, a temporary readvance of the Cape Cod Bay lobe led to the formation of the Sandwich moraine by the tectonization of the sediments in front of the lobe. A similar readvance, but somewhat earlier, by the Buzzards Bay lobe also deformed the sediments along the western edge of the Mashpee Pitted plain to form the Buzzards Bay moraine. Shortly after 18,000 B.P., the Cape Cod Bay lobe retreated farther north to a position in the center of Cape Cod Bay and the Buzzards Bay lobe northwest of Buzzards Bay. The South Channel lobe not only retreated northward, it also retreated eastward, creating an intervening low between itself and the Cape Cod Bay lobe. As the South Channel lobe retreated away from the northern edge of Georges Bank and the head of Great South Channel, marine waters invaded the Gulf of Maine

26

E. Uchupi and Others

depression via Northeast Channel with a sill depth of 232 m (Fig. 21). These waters flooded westward along the base of the northern slope of Georges Bank reaching the vicinity of Great South Channel. A proglacial lake, Cape Cod Bay Glacial Lake, also was formed at this time between the retreating Cape Cod Bay lobe, the drift deposits in the upper Cape and South Channel lobe to the east. As the ice retreated, the lake in Cape Cod Bay south of it grew northward, keeping pace with the retreating ice lobe until it covered an area of more than 1,000 km2. The fine-grained sediment beneath the Provincetown Hook

sands may represent the final stages of deposition in this lake, as it drained soon after the low between the Great South and Cape Cod Bay lobes reached the vicinity of Provincetown. The lake drained first via Manomet Valley to Buzzards Bay and later via Bass River to Nantucket Sound and finally via Town Cove to the low at the Cape’s elbow. Deposition in the lake was via subglacial channels along the grounded Cape Cod lobe front to the north and from the South Channel lobe to the east. Some sediment also entered the lake from the south (as attested by the same deltas), north of the Sandwich moraine.

Figure 19. Map showing the southern limit of the Laurentide ice sheet, the distribution of the subglacial channels in Nantucket Sound, and the southern limit of abundant gravel. Compiled from Schlee and Pratt (1970, Fig. 15) and Oldale (1982, Fig. 3).

Late Quaternary construction of Cape Cod, Massachusetts The plains of the lower Cape, like those of the upper Cape, also were deposited by a braided stream system, but in this case by a system that originated in the South Channel lobe east of Cape Cod. During the formation of the lower Cape plains, the western ice front of the South Channel lobe was located about 3 to 7 km east of the present coast of the lower Cape. The first plain to be formed was Wellfleet plain, followed by the Highland, Truro, and Eastham plains (Fig. 2). Prior to the deposition

27

of the Wellfleet plain, the Cape Cod Bay lobe readvanced southward, deforming the glaciolacustrine sediments in front of it to form Billingsgate Shoal. The Buzzards Bay lobe also readvanced at the same time to form the Monks Hill moraine along the west coast of Cape Cod Bay (Fig. 21). Sediments of the Highland plain (Fig. 2) may have been deposited in a temporary lake formed as a result of partial collapse of the outwash plain along the western margin of Great South Channel. The

Figure 20. Paleogeographic map of the Cape Cod region just prior to 18,000 B.P. Map shows positions of the Buzzards Bay, Cape Cod Bay, and South Channel ice lobes at that time and the extent of a possible glacial lake in the present area of Nantucket Sound. Map is based on data discussed in the text. The valleys in Martha’s Vineyard and Nantucket Island were eroded at this time by springs that originated in the lake that occupied the present position of Nantucket Sound. Possibly the same process also is responsible for the initial erosion of Muskeget Channel between Martha’s Vineyard and Nantucket Island.

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E. Uchupi and Others

Figure 21. Paleogeographic map of the Cape Cod region soon after 18,000 B.P. Location of glacial front and associated glacial lakes is from Larson (1982, Fig. 3), and the valley distribution from O’Hara and Oldale (1987, Fig. 6) and Uchupi and Oldale (1994, Figs. 1, 2). Larson (1982) estimated that the Cape Cod Bay Glacial Lake ultimately covered an area of more than 1,000 km2 before it drained.

lake was initially formed after deposition of the Wellfleet plain and grew as the proximal end of the Truro plain collapsed after its deposition. The Eastham plain was formed from runoff from a sublobe of the Great Channel lobe that occupied the present position of the marshes and bays west of Nauset Beach and the filled low at the Cape’s elbow. This low extended across the Cape connecting Cape Cod Bay and the Gulf of Maine via Town Cove. Seeps from the proglacial lake in Cape Cod Bay and the lake atop the outwash plain east of the lower Cape eroded the

channels in Cape Cod. One of the valleys in the lower Cape, Pamet River, may have extended across the lower Cape to the lake, leading to the catastrophic drainage of the lake westward into Cape Cod Bay (Fig. 2). The channels in the upper Cape may have extended their courses into Nantucket Sound and beyond, forming the drainage system eroded out of the Pleistocene sediments in the region (Fig. 21). The valleys in the sound have reliefs of up to 14 m and thalweg depths of more than 34 m, with a dendritic pattern indicating drainage to the east to Great South Channel, south toward Muskeget Channel

Late Quaternary construction of Cape Cod, Massachusetts between Chappaquiddick and Muskeget Islands, and southwest toward Vineyard Sound between the Elizabeth Islands and Martha’s Vineyard (Figs. 1, 21). In Vineyard Sound, the valleys have local relief up to 20 m, thalweg depths greater than 55 m, and a dendritic pattern indicating drainage southwestward. Once the lakes in Cape Cod Bay and east of the lower Cape drained and the groundwater table level dropped, valley formation on the Cape and vicinity ceased. By 17,600 B.P., marine conditions were well established in the western Gulf of Maine, conditions that by no later than 16,000 B.P. had extended to the region of Stellwagen Basin as a

29

result of crustal subsidence by ice loading (Fig. 22). At this time meltwater deposition from the South Channel lobe formed Stellwagen Bank in the wide reentrant between the South Channel and Cape Cod Bay ice lobes. Glaciomarine sediments in Stellwagen Basin and the region south of the basin were derived mainly from the Cape Cod Bay lobe. From this basin marine waters migrated southward to the northern rim of the area formerly occupied by Cape Cod Bay Glacial Lake (i.e., the southern edge of Cape Cod Bay). About 3,000 yr later (about 14,000 B.P.), northeastern Massachusetts was ice free and covered by marine waters as a result of its depression by ice loading. As the

Figure 22. Paleogeographic map of the Cape Cod region about 16,000 B.P. Compiled from various sources discussed in the text.

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E. Uchupi and Others

crust rebounded isostatically about 12,000 B.P. (2,000 yr after deglaciation) after its ice load was removed, sea level dropped from an elevation of +33 m above the present level to at least –43 m to possibly –55 m below its present level (Fig. 23). The north-draining fluvial system in Cape Cod Bay was cut out of the proglacial and glaciomarine sediments during this regression. Valleys in this dendritic system have reliefs exceeding 20 m and thalweg depths exceeding 70 m. The amorphous sand above the glaciomarine sediments may represent a deltaic facies deposited by this drainage system, or it could be older and may

be a debris flow emplaced prior to the regression (Fig. 22). From this depth sea level again began to rise, reaching to near its present level about 1,000 B.P. As the sea transgressed across the exposed shelf, it gradually drowned the preexisting fluvial-modified glacial and glacial terrains. The sea first flooded the valleys carved out of the shelf strata, creating a complex estuarine system. By 10,000 B.P., when sea level was at a depth of –50 m, the mouth of the channel in Vineyard Sound was flooded as well as those of the valleys in Cape Cod Bay. When sea level rose to –30 m about 9,000 B.P., Muskeget Channel flooded. As sea level

Figure 23. Paleogeographic map of the Cape Cod region about 12,000 B.P. during the latest Wisconsinan regression. Distribution of fluvial valleys in Cape Cod Bay compiled from Oldale and O’Hara (1990, Fig. 8).

Late Quaternary construction of Cape Cod, Massachusetts continued to rise, the valleys in Nantucket Sound were filled with silts and clay of nonmarine and estuarine origin, and by 6,000 B.P. only a narrow coastal zone remained exposed, which ultimately was drowned when sea level approached to near its present level 5,000 yr later (Fig. 24). ORIGINAL GLACIAL SHAPE OF CAPE COD Toward the end of the Younger Dryas (10,500 to 10,000 B.P.) and before the initiation of the second meltwater pulse centered at 9,500 B.P. when sea level was about –30 m below

31

its present level, glacial Cape Cod was much larger than it is today having had an area of 1,550 km2. During the subsequent rise in sea level, this area was reduced to 1,075 km2, 69% of the original glacial Cape Cod (Table 1, Fig. 2). Since then, about 140 km2 have accreted onto this glacial core by marine processes. The southern margin of the glacial Cape was approximately 1 to 3 km south of the present coast and was marked by a southerly facing slope having a relief of about 10 m and a gradient of about 1° to 2°. The southerly draining valleys formed by spring sapping probably originated in this free surface and from there propagated northward in the manner

Figure 24. Paleogeographic map of the Cape Cod region about 6,000 B.P. Compiled from various sources discussed in text.

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E. Uchupi and Others TABLE 1. AREAS OF GLACIAL GEOMORPHIC TERRAINS OF CAPE COD % of Glacial % of Present Cape Cape

Terrain

Area (km2)

Glacial Cape Cod Glacial remnant Eastern terrace (eroded) North terrace (drowned and eroded) Billingsgate Shoal Coastal segments

1,562.3 1,077.4 193.2 11.4 182.6 97.9

0.7 11.7 6.3

Present Cape Glacial remnant Provincetown Hook Barnstable Marsh, Nauset Spit Nauset Beach and adjacent marshes Monomoy Island Beaches, bars, spits North shore South shore West shore

1,217.5 1,077.4 34.6

77.9

100.00 69.0 12.4

100.00 88.5 2.8

12.0

1.0

21.8 9.5

1.0 0.8

38.1 11.8 12.5

3.1 1.0 1.0

described by Uchupi and Oldale (1994). Drainage down the free surface to Nantucket Sound also carved a fluvial system in this region, with Vineyard Sound and Muskeget Channel serving as the main water gaps of the system (Fig. 21). The glacial Cape’s northern margin was 1 to 0.5 km north of the present shore and its west side about 1 to 1.5 km west of the present one. These boundaries also were marked by low cliffs having reliefs on the order of 10 m and gradients of about 1° to 2°. Like valleys in the upper Cape, those on the lower Cape also originated along this free surface and from there propagated upslope eastward and extended their courses westward downslope onto the lowland beyond Billingsgate Shoal (Fig. 2). Extending southwesterly from the west side of the glacial Cape for a distance of 18 km was Billingsgate Shoal, a broad moraine having a relief of about 10 m. On its northern side, glacial Cape Cod terminated on a north-facing cliff with a relief of nearly 80 m and a gradient of 2° (Fig. 25D). Our investigation indicates that the lower Cape was much wider than Davis (1896, 1954) believed when he proposed that the glacial lower Cape extended about 4 km east of its present position. Our study shows that at 70°5′W, the glacial Cape coast was 2.8 km east of the present one; at 70°W 4.5 km east, just slightly north of 42°N, 5.5 km; 6.3 km east at 42°N, 7.0 km at 39°55′N; and south of 39°55′N it was 6.0 km east of the present coast. Like the present coast, this glacial shore also was characterized by an eastward-facing scarp. The cliff’s relief when sea level was –40 m below the present one ranged from 95 m at about 39°57′N to 60 m just north of 39°50′N and 70°05′W. Today the relief of the eastward-facing Atlantic cliffs of the lower Cape ranges from over 40 to 18 m above the present sea level with a maximum gradient of 35°. Of the spring sapping valleys on the

lower Cape, only Pamet Valley probably extended some distance eastward of the present coast, terminating on the possible lake that was formed atop the clastic wedge deposited by the South Channel lobe. Throughout much of the Cape’s elbow, the edge of the glacial Cape was 1 to 4 km west of the present shore. Only at the northern end of Nauset Beach was glacial Cape Cod east of the present Cape, at a distance of about 1 km east of its present position. In this segment of glacial Cape Cod was a westerly trending nearly filled embayment with a relief of 10 to 40 m (Fig. 25E). Although the southern and northwestern margins of glacial Cape Cod appear to be irregular, the rest of the Cape’s upland periphery may have been smoother. As a whole, the Cape does not appear to have had the irregular outline postulated by Davis. Topographic irregularities appeared to have been restricted to the plains where spring sapping valleys and kettle holes produced a hummocky topographic texture on the plains, or to ice contacts. The glacial Cape upland was essentially two sediment plains with the one in the upper Cape dipping southward in the direction of Nantucket Sound, and the one in the lower Cape dipping westward in the direction of Cape Cod Bay. These two sediment accumulations were separated by a broad embayment that may have served as a connection with the Gulf of Maine and Cape Cod Bay lowland. South and north of the upland were low areas currently occupied by Nantucket Sound and Cape Cod Bay. East of the upland was a marine embayment that extended the length of the Wilkinson Basin complex in the western Gulf of Maine. Present positions of Provincetown Hook and Stellwagen Basin also were inundated by these marine waters, with only Stellwagen Bank rising above its general level to form an island. HOLOCENE MODIFICATIONS Phase I: 9,500–6,000 B.P. Modifications of glacial Cape Cod by marine processes during the Holocene transgression can be divided into two phases: Phase I, which began about 9,500 B.P. and terminated about 6,000 B.P.; and Phase II, from 6,000 B.P. to the present. During Phase I, sea level rose at a rate of about 6 m/1,000 yr from –30 to –10 m below its present level. In Phase II sea level rose to near its present level about 1,000 B.P. at a rate of 2.0 m/1,000 yr, or about one-third of the previous rate. During this period crustal uplift due to ice unloading probably terminated and the crust began to subside again as the peripheral bulge decayed. This movement may still be taking place today accounting in part for a relative rise in sea level of about 2 to 3 mm a–1. The major modifications of glacial Cape Cod during Phase I were massive erosion of the east and north sides of the lower Cape, the filling of the low at the Cape’s elbow and the drowning of the Billingsgate Shoal moraine, Stellwagen Bank, most of Nantucket Shoals and Sound, and Cape Cod Bay. During the flooding of Nantucket Shoals and Sound, tidal currents and

Late Quaternary construction of Cape Cod, Massachusetts

33

Figure 25. Schematic geologic cross sections (locations shown in Fig. 2) of the lower Cape. C, The section extending the length of the Provincetown Hook was compiled from data of Zeigler et al. (1965, Fig. 2), and Oldale (1988, Fig. 5). This cross section differs from the one in Zeigler et al. (1965) in that we assumed that the red clay in the well at Pilgrim Lake is a postglacial deposit and not part of the Truro plain as postulated by them. D, The section extending eastward from the lower Cape was reconstructed from the seismic reflection profile in Figure 14 and geophysical data from the lower Cape (Oldale, 1969) and Cape Cod Bay (Oldale and O’Hara, 1990). E, The section in the area of the Cape’s elbow was compiled from the seismic reflection profile in Figure 13, and geophysical data from the Cape (Oldale, 1969) and Cape Cod Bay (Oldale and O’Hara, 1990).

waves led to extensive reworking of the sediments, and the former outwash plain southeast of Cape Cod was slowly transformed to the present ridge-trough configuration of Nantucket Shoals. Most of the detritus winnowed from the sediments was transported westward to form Fan A and into the Mud Patch along the shelf’s edge south of Nantucket Island. Although poorly documented, the northwest-facing cliff defining the northwest tip of the Truro plain apparently retreated about 3 km prior to the formation of Provincetown Hook and a relatively steeply dipping terrace with a gradient of 0.8° was cut out of the cliff (Fig. 25C). Possibly this bench-like feature is sediment block that slumped from the cliffs as they retreated eastward. The debris-like sequence in the subsurface at Race Point and Provincetown may have originated in this slump (Fig. 25C). Cliff retreat along the western edge of the Truro plain and the east side of the lower Cape was due to wave

action and submarine erosion to a depth of about 10 m. This profile of erosion was thought to be in equilibrium with longterm sea state (Zeigler et al., 1964b). Thus, the outer edge of the terrace on the east side of the lower Cape at a depth of –40 m was probably cut at the time when sea level was at a depth of –30 m about 9,500 B.P. Submarine erosion was probably as important in producing the erosional surface as were processes at sea level. Studies by Zeigler et al. (1964a, Table 1) on the eastern lower Cape indicate that erosion below sea level today is slightly greater than wave erosion at the coast, yielding a yearly average volume of about 330,000 m3 of sediment compared to the yearly sediment volume provided by coastal wave erosion of 310,000 m3, a combined total of 640,000 m3. Cliff retreat by wave erosion is the result of undercutting creating oversteepening, subsequent collapse of a segment of the cliff, and transport of the slumped mass by waves. When the

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E. Uchupi and Others

cliff was initially bordered by deep water, the slumped sediment would tend to accumulate beyond the cliff and the cliff would be reexposed to erosion immediately, but as the terrace widened and the slump mass accumulated at the foot of the cliff, it was protected from further erosion by wave action until the slumped sediments were dispersed. As collapse was a local phenomenon, cliff retreat probably occurred in segments. A cliff segment would be undercut by wave action and a slump would take place. This segment of the cliff would then be temporarily protected from further erosion, and erosion would shift to another cliff segment. While this segment was being undercut and another slump was formed, the older slumped mass would be reworked by waves again exposing the cliff segment to wave erosion. These undercutting/slumping/transport events along the cliff face are so close in time, however, that within the resolution of geologic time cliff retreat is synchronous along its length. The eastward-facing sea cliffs of the lower Cape are retreating today at the rate of about 0.8 m a–1 (Zeigler et al., 1964b). Thus, 1,000 B.P., when sea level reached its present position, the cliffs were 800 m east of their present position. The seaward edge of the terrace fronting the cliffs at a depth of about –40 m marks the position of the cliffs at about 9,500 B.P., when sea level was at a depth of –30 m. The rate of cliff erosion was greater at the point where the terrace fronting it is widest, as attested by the values given below. The average rate of cliff retreat from 9,500 to 1,000 B.P. ranged from 0.3 to 0.5 m a–1 north of 42°N where the terrace is 2,700 to 4,000 m wide, 0.6 to 0.9 m a–1 from 42°N to 41°53′N where it is 5,000 to 8,000 m wide, and 0.5 m a–1 south of 41°53′N where the terrace is 4,000 m wide. The surface of this wave-cut terrace, however, does not display a uniform gradient and its continuity is disrupted by three groups of secondary terraces with declivities of 7′ or less. The slopes separating these secondary terraces display gradients ranging from 1°13′ to 9′ (Fig. 26A). These slope variations are particularly noticeable in depths shallower than about 30 m. As the declivities of these erosional surfaces were controlled by rise in sea level and rate of erosion (i.e., gradient-rate of sea level rise/rate of erosion or cliff retreat), then either the rate of sea-level rise or rate of erosion changed as the terrace was cut. That there may have been episodic minor stillstands during the Holocene submergence of southern New England is suggested by paleoshoreline features in Long Island Sound at depths of –37, –27, and –25 below present sea level (Gayes and Bokuniewicz, 1991). Inferred ages of these shorelines range from 10,500/10,000 B.P. for the –37-m terrace to 9,500/9,000 B.P. for the –25-m structure. As the gradients off the lower Cape varied not only in a seaward direction but also along strike, then these factors (sea-level rise, rate of cliff retreat) must have varied both with time and place to account for gradient variations along strike. The only way the relative sea-level rise could vary along the 33-km length of the terrace is if tectonism due to glacial unloading, collapse of the peripheral bulge, or some other process varied along the terrace’s strike. Groen et al. (1986) and

Figure 26. A, 3.5-kHz echo-sounding profiles of the erosional terrace east of the lower Cape. Inset map shows locations of the profiles. Terraces along the profiles are indicated by bracketed heavy line. Numbers along the edge of the terrace (dotted line) in the inset map represent the relief of the clastic wedge before the terrace was eroded during the Holocene transgression. This relief is relative to the depth of sea level at about 10,000 B.P. when it was –40 m below its present level. The numbers along the coast are the relief (in meters) of the present cliff above sea level. B, Depth distribution of terraces along profiles.

Stone and Borns (1986), for example, suggested that isostatic rebound in southern New England was between 0.7 to 0.9 m km–1 to the north of Cape Cod. However, the sporadic variations of gradients in that direction fail to support either isostatic rebound or collapse of the peripheral bulge as the causes for the change in gradient along strike. Thus, this change in the terrace’s gradient along strike is probably due to variations in the rate of cliff retreat. Rate of erosion is controlled by oceanographic factors and the erodibility of the sediments being removed. This erodibility is controlled by the texture and fabric of the sediments. The lithology of the present cliffs consisting of cross-bedded sands, gravelly sands, and a clay layer, changes unpredictably along the strike (Fig. 28; Zeigler et al., 1964b; Koteff et al., 1967). The sections measured by Zeigler et al. (1964b) and Koteff et al. (1967) were at the distal end of the

Late Quaternary construction of Cape Cod, Massachusetts sediment wedges deposited by the South Channel lobe where sediment variations along the strike would be minimal. Such lateral changes in facies were probably even more pronounced near the sediment sources. Yet as shown by Figure 26B, the outer edge of the terrace displays a relatively uniform gradient along the strike; it is here that the secondary terraces are best developed, suggesting that cliff retreat at that time was relatively uniform along the strike. Variations in gradients along the terrace’s surface also are difficult to explain as the result of oceanographic factors. As is the case today, maximum erosion of the cliffs probably was by northeasterly storms. Although these storms varied in their intensity with time, there is no obstacle to preventing one segment of the east-facing cliffs from being attacked more intensely than another. Although the crest of Georges Bank was above sea level at the time terrace cutting was initiated 9,500 B.P., and may have acted as a barrier from long-period waves from the east and southeast as noted by Zeigler et al. (1965), this crest would not explain why one part of the cliffs was eroding faster than another. Thus, at present we are not able to separate the causative sources responsible for the formation of the erosional terrace. Whatever the causes of this variation in the rate of erosion along strike, they must even out over time as the face of the present cliff is relatively smooth, an outline also displayed in older charts (Zeigler et al., 1964b). The terrace eroded from 9,500 to 6,000 B.P. (the position of the 6,000 B.P. shoreline was determined by extrapolation from the positions of the 1,000 and 9,500 B.P. shorelines) has an area of 86 km2. The sediment wedge eroded to form the erosional terrace ranged in thickness from 95 to 60 m along its seaward edge to 70 to 45 m at its landward edge, averaging about 70 m in thickness. Thus, on the order of 6.1 km3 of sediment were eroded both above and below sea level to form the terrace. Today about 82% of the glacial material eroded from the cliffs is available to nourish the coastal features with 18% being lost to deeper water offshore (Zeigler et al., 1964b, 1965). If similar partitioning existed 9,500 yr ago, about 1.1 km3 eroded from the cliffs was transported eastward to form Fan C. Contribution to the fan probably exceeded this amount when the shoreline was located closer to the terrace’s edge. Slope retreat then as today (Figs. 27, 28) must have been enhanced by mass wasting (slumping and gravitational sliding), as evidenced by the debris flow at the base of the slope (Fig. 14). As the cliff retreated westward, sediment contribution to Fan C slowly diminished and most of the finer sediments were probably transported eastward beyond the fan into deeper water. The remainder of the sediments eroded from the cliffs, like today, initially accumulated in beaches and nearshore (43.3%) and offshore bars (36.3%) (Zeigler et al., 1964a, Table III). The gravel fraction tended to concentrate at the toe of the beach, although it may have spread over the foreshore part of the beach or part of the inner bar during some sea conditions (Zeigler et al., 1964a). In addition to gravel, the beach and nearshore bar also tended to be sites of very coarse to coarse sand deposition and the finer fraction accumulated in the offshore bar (Fig. 29). On the assumption that the

35

Figure 27. The eastward-facing sea cliffs in the area of the Highland Light in Truro. Note the two massive relict slump structures with their extensive vegetation cover in the center of the photograph. The smaller gravitational structures defacing the cliff do not have this cover attesting to their more recent age. The shoreline also shows repetitive hooked bars attesting to ample sediment supply. Photograph by D. G. Aubrey, courtesy of the Public Information Office, Woods Hole Oceanographic Institution.

coastal features are neither growing nor diminishing in volume, Zeigler et al. (1964a) estimated that the residence time of the sediments in the beaches and nearshore bars was 38.2 yr and on the offshore bar 57.2 yr. Similar residence times probably also characterize the formation of the terrace. At that time, 9,500 B.P., Georges Bank was partially exposed, a barrier limiting long-period waves from the east and southeast as a result of which sediment transport to the north was inhibited. However, no such barrier existed for waves from northerly directions at that time, and as a result, sediment transport to the south was enhanced (Zeigler et al., 1965). The depocenter of this southward transport was the nearly filled offshore low at the Cape’s elbow including the region of the pre-

36 E. Uchupi and Others

Figure 28. Diagrammatic stratigraphic sections from the Highland Light Tower region in the lower Cape. The numbers at the top of the figure indicate distances in meters between the sections. From Koteff et al. (1967).

Late Quaternary construction of Cape Cod, Massachusetts

Figure 29. Textural variation of the drift on the cliff and the sands on the beach, and near- and offshore bars on the eastern lower Cape. Modified from Zeigler et al. (1964a, Fig. 6). Note that the larger medium diameters occur in the beach and nearshore bar (0.78 to 0.74 mm) and the finest in the offshore bar (0.35 mm). The drift on the cliff has an intermediate size of 0.58 mm.

sent marshes and bays west of Nauset-Chatham beaches. The marine sediments filling these lows have an area of about 260 km2, and range in thickness from 45 to 60 m between Pleasant Bay at Strong Point and Chatham Bars Inlet (LeClair et al., 1978), 10 to 30 m on the shelf (Aubrey et al., 1982), and 40 to 20 m from the outer shelf to the base of the slope. Their average thickness is about 30 m, forming a volume of 7.8 km3. About 64% (5.0 km3) of this fill was derived from the north, with the remaining 36% coming from the Harwich and Eastham plains west of the depression. The sediments near the former Chatham Bars Inlet rest on an irregular reflector that LeClair et al. (1978) and McClennen (1979) suggested may be the erosional surface of Tertiary(?) age or a boundary within the Pleistocene. We suspect that it marks the top of the Wisconsinan glacial deposits and is a transgressive unconformity. The thickness of the postglacial unit in the Chatham Bars Inlet may indicate either a nearby source or that LeClair et al. (1978) used an unusually high velocity to convert travel time to thickness in meters. The upper 6 m of the Holocene marine sequence near Chatham Bars Inlet are disrupted by filled channels that are similar in size to the present tidal inlet near Morris Island (LeClair et al., 1978; McClennen, 1979). On the shelf the unit is acoustically diverse lacking internal reflectors in places and in others having many discontinuous horizons (Aubrey et al., 1982). Lack of unconformities within the upper unit indicates that its emplacement was continuous. As near Chatham Bars Inlet, the Holocene sediments on the shelf

37

are separated from the more massive glacial lower unit deposited by a sublobe of the South Channel lobe by a prominent reflector. This surface is channelized, features which Aubrey et al. (1982) believed were eroded subaerially by discharge from Cape Cod Bay Glacial Lake. Along the shelf’s edge the unit is characterized by a sequence of foreset beds dipping toward the east and south. These trends point to both northerly and westerly sources for the Holocene sediments in the region. Eastward progradation toward the Gulf of Maine led to the construction of Fan B and westward progradation partially filled the lows west of Nauset-Chatham beaches. Once they were filled and the barrier beaches along the present shore developed, the coastal lows become sites of marsh accumulation. During Phase I, sediments derived from erosion from the northwest-facing cliffs defining the western limit of the Truro plain also were transported southward by littoral drift. The depocenter of these sediments, however, was in Cape Cod Bay west of the lower Cape. During Phase I, extending from 9,500 to 6,000 B.P., Stellwagen Bank and most of Billingsgate Shoal moraine, Cape Cod Bay, Nantucket Shoals, and Nantucket Sound were drowned. As Stellwagen Bank disappeared beneath the sea, its top was truncated with the resulting detritus being transported both westward into Stellwagen Basin and southward into Cape Cod Bay. During the erosion of the bank a lag deposit of very coarse gravel was formed on its top, a cover that has served as an armor protecting the bank top from further extensive erosion. Also during Phase I, the hook at the southwestern end of the Stellwagen Bank may have formed from sand displaced from the top of the bank. Construction of the spit terminated about 7,500 B.P. when the 18-m-deep northwest part of the bank was submerged (Zeigler et al., 1965). As the Billingsgate Shoal moraine was submerged below sea level, its crest also was smoothed by a combination of erosion of the highs and filling of the lows. The southwesterly tip of the shoal was truncated at this time, and the resulting erosional surface was buried by sediments derived from the northeast- and the northwest-side slopes of the shoal. These sediments were transported to their site of accumulation by southwestward littoral drift. As the marine waters transgressed over most of Cape Cod Bay and Nantucket Sound, an extensive unconformity was cut on the sediments. Decelerations of the transgression led to local accumulation of beach and bar deposits on top of the unconformity. Phase II: 6,000 B.P.–Present North, south, and west shores and Sandy Neck/Barnstable Marsh. From 6,000 to 1,000 yr ago when sea level approached to near its present level, the rate of relative sea-level rise decreased from 6 to 2 m ka–1, and crustal subsidence due to the collapse of the peripheral bulge, which began during Phase I, gradually died out or slowed down. The relative rise in sea level of 2 to 3 mm a–1 documented by tide-gauge data during the last 100 yr may result in part from other tectonic forces.

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During this period of relative low rate of sea-level rise, features such as Sandy Neck/Barnstable Marsh, Provincetown Hook, Nauset Beach and associated marshes, and Monomoy Island were built (Figs. 30 through 33). Also at this time, the drowning of Billingsgate Shoal, Cape Cod Bay, and Nantucket Sound and Shoals was completed, and the Cape’s south, north, and west shores retreated 1 to 3 km from their former positions (Fig. 2). The numerous bars and pocket beaches, and small spits fronting the glacial terrain in the south, north, and west shores began to form after this retreat, a development that still continues today. On the south shore these constructional forms have a composite area of 11.8 km2, making up 0.97% of the area of the present Cape (Table 1). Similar features, exclusive of Sandy Neck Spit, on the north shore have an area of 36.1 km2 making up 3.13% of the present Cape, and on the west shore they have an area of 12.5 km2 constituting 1.03% of the present Cape. The littoral features on the south shore were formed by both easterly and westerly flowing drift, those on the north shore mostly by an easterly flowing one, and the ones on the west coast by both a south- and north-flowing littoral drift. As the south shore retreated northward, the south-trending spring sapping valleys in the upper Cape were flooded to form the elongated bays that dominate this shore today (Fig. 34). Connection of coastal kettle lakes via tidal inlets, such as Oyster Pond, has added further irregularities to the outline of the south shore of the upper Cape (Fig. 35). The most prominent constructional feature of the upper Cape is Sandy Neck Spit and associated Barnstable Marsh, having a combined area of 12.0 km2 (the spit has an area of 10.5 km2 and the marsh, 1.5 km2); they make up 0.98% (0.86% of which represents the spit and the remaining 0.12%, the marsh) of the present Cape (Table 1). Redfield (1965) recon-

structed from soundings and borings of the peat the development during the past 4,000 yr of the salt marshes and Sandy Neck Spit that shelters them; its rate of vertical accretion and rate of sea-level rise were determined from 14C age determinations of the peat. During the creation of the spit-marsh complex, sea level rose at a rate of about 3 × 10–3 m a–1 from at least 3,700 to 2,100 B.P. and about 1.0 × 10–3 m a–1 during the last 2,100 yr (Redfield and Rubin, 1962). Interpretation of the vertical growth of the marsh was determined from the vertical distribution of high marsh peat (Spartina patens, a dwarf form of Spartina alterniflora, and Distichlis spicata), which flourishes at high water level, and the intertidal peat (Spartina alterniflora) extending from high water level to nearly two-thirds of the tidal range. The earliest stage reconstructed by Redfield was at 3,793 B.P. when sea level was –5.5 m below its present level. At that time Sandy Neck was less than 2 km long (Fig. 36A). With the extension of the spit and sediment accumulation behind it, the marsh became more continuous in the enclosed basin (Fig. 36B). Further growth of the marsh (Fig. 36C–E) appears to have been in the form of tongues that occupied positions where sand flats had developed. As the vegetation on these flats grew, the water in the enclosed basin was first separated into broad sounds, and as the vegetation continued to grow, the sounds became narrower to form the present creek systems. As the peat built up to the present high water line, the vegetation stabilized the channel meander pattern. Eastward growth rate of the spit appears to have diminished since about 2,000 yr ago as a result of a decrease in the rate of sea-level rise at that time. Redfield (1965) postulated that the terminal hooks at the end of the spit (Fig. 36D–F) developed during this period of slowly rising sea level. As summarized by Redfield, the ontogeny of Barnstable

Figure 30. Photograph of Barnstable Marsh looking east with Sandy Neck on the left and the marsh on the right. Photograph courtesy of D. S. Blackwood, U.S. Geological Survey, Woods Hole.

Late Quaternary construction of Cape Cod, Massachusetts

Figure 31. Photograph of Provincetown Hook looking east with High Head sea cliffs in the background. Pilgrim Lake is on the left and on the upper right side is Wellfleet Harbor. To the left of Pilgrim Lake are parabolic dunes whose beach grass cover was planted in an attempt to stabilize the dunes. Photograph courtesy of D. S. Blackwood, U.S. Geological Survey, Woods Hole.

Figure 32. Photograph of Nauset Spit and marsh looking south. Photograph courtesy of D. S. Blackwood, U.S. Geological Survey, Woods Hole.

39

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Figure 33. Photograph of Nauset Beach and North and South Monomoy Island and adjacent mainland looking toward the south. At the time this photograph was taken in November 1980, the Chatham Break at the position indicated by the arrow did not exist. Compare with Figure 52. Note also the 1978 break in Monomoy Island. Photograph by A. G. Gaines, Jr., courtesy of the Information Office, Woods Hole Oceanographic Institution.

Figure 34. Photograph of drowned spring sapping valleys and kettle holes in Falmouth in the upper Cape. Photograph courtesy of B. Howes, Biology Department, Woods Hole Oceanographic Institution.

Late Quaternary construction of Cape Cod, Massachusetts

41

Figure 35. Photograph of Oyster Pond at Falmouth in the upper Cape. This kettle lake is connected to Vineyard Sound via a tidal inlet located in the left side of the photograph. Photograph courtesy of B. Howes, Biology Department, Woods Hole Oceanographic Institution.

Marsh is the result of the eastward growth of Sandy Neck during the last 3,800 yr and the spread of the marsh in the basin enclosed by the spit. As a result of higher sedimentation rates, marsh expansion was much faster along the upland than along the inward margin of the spit, and the marsh advanced basinward in the form of broad tongues resting on sand flats. As the marsh continued to grow, the sounds between the vegetation tongues were transformed to today’s creeks and meandering channels, features that change slowly only under existing tide and sea-level conditions. East shore, Provincetown Hook to Monomoy Island. At 6,000 B.P. when sea level was –10 m below its present level, the east shore of the lower Cape was still dominated by the 32-km-long east-facing cliffs located 4 to 2.5 km east of their present locations. At that time, the present site of Provincetown Hook north of the cliffs was a marine embayment connecting the western Gulf of Maine and Cape Cod Bay. The southeast end of the Cape also was the site of a marine embayment, a low that also may have been connected to Cape Cod Bay via Town Cove. Georges Bank was finally submerged by 6,000 B.P. and no longer inhibited waves from the east and southeast. As a result of this drowning, northward littoral currents became more pronounced and began the construction of Provincetown Hook (Zeigler et al., 1965). From 6,000 to 1,000 yr ago the sea cliffs in the lower Cape retreated distances of 1.5 km at 70°05′W, 2.0 km at 70°W, 2.4 km at 42°N, and 3.0 to 3.2 km farther south at rates ranging from 0.3 m a–1 at the northern end

to 0.6 m a–1 at the southern end of the terrace. As the cliffs retreated westward, they encountered the relict spring sapping valleys. From 1887 to 1957–1958 this rate increased to 0.8 m a–1 (Zeigler et al., 1964b). The causes of this increase have yet to be resolved, but the fact that relative sea-level rise also accelerated during that period suggests that the acceleration may be responsible for the increased erosion rate. During the coastal retreat a wedge of glacial Cape Cod with a thickness of 45 to 70 m on the seaward edge and 25 to 50 m on the landward edge, having an area of approximately 78.9 km2, was destroyed; a volume of 3.6 km3 (assuming an average thickness 45 m above the erosional terrace) of detritus was removed from the site and transported elsewhere. During the last 1,000 yr when the cliffs retreated to their present position, another 14.8 km2 were eroded from the lower Cape. Assuming an average thickness of 30 m this equates to a volume of 0.4 km3 for a total volume of 4.0 km3 for the last 6,000 yr. This equates to sediment production of 666,667 m3 per year. According to Zeigler et al. (1964a) (Table 1), the average yearly volume today is 644,428 m3, of which 526,000 m3 (82%) was available to nourish the beaches, bars, and spits. In our calculations this amounts to 546,667 m3, which converts to a volume of 3.28 km3 for the last 6,000 yr. The segment of the Provincetown Hook above sea level has an area of 34.6 km2 and an average thickness of Holocene sediments of 20 m, for a volume of 0.7 km3. The segment of the hook below sea level has an area of 37.6 km2 and an average sediment thickness of about 60 m for

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Figure 36. Paleogeography of the Barnstable estuary and marsh. Modified from Redfield (1965). Numbers at top right, A through F, indicate the age of the map and the mean high water level (in meters) below present level at that time. Although Redfield assigned a date of 1950 A.D. to the map in panel F, we have assigned a date of 1993 to this map on the assumption that the changes that have occurred during the past 43 yr are beyond the resolution of the illustration.

a volume of 2.26 km3, or a total volume of the hook of about 2.96 km3. Zeigler et al. (1965, Table 3) calculated the volume of the Provincetown Hook as 2.8 km3, a volume similar to that determined in the present study. They also postulated that 0.08 km3 (a yearly production of 15,299 m3) of this sediment originated from the Cape Cod Bay side of the Cape during the last 6,000 yr, adding to Provincetown Hook. The remainder, some 2.7 km3, came from erosion of the cliffs south of the Provincetown Hook. This leaves about 0.45 km3 of sand for the construction of the bars and beaches fronting the cliffs and those south of the cliffs. Zeigler et al. (1965, Table 3) calculated that the volumes of the beaches, nearshore and offshore bars is 0.004, 0.006, and 0.013 km3, respectively, for a combined volume of 0.024 km3, leaving 0.21 km3 for the construction of the features farther south. According to Zeigler et al. (1965), the volumes of the fea-

tures south of the cliffs were as follows: Nauset Spit, 0.007 km3; Nauset Beach, 0.05 km3; Monomoy Island, above and below sea level, 0.23 km3; Handkerchief Shoals, 0.14 km3; Great Round Shoals, 0.25 km3; Stonehorse and Little Round Shoals, 0.28 km3; and Bearse Shoals, 0.07 km3, for a total of 1.01 km3. The east-west orientation of the shoals and their isolation by Pollack Rip and Great Round Channels where the glacial sediments (postglacial transgressive unconformity) are exposed suggest that the sediments making up the Great Round, Stonehorse, and Little Round Shoals were not derived from the north (Fig. 37). Their grain size (1.0 to 0.250 mm on the lows between the shoals and 0.250 to 0.0625 mm in the shoals themselves; Schlee, 1973, Plate 2E) also suggests that these sediments were probably derived from Nantucket Sound with textural variations reflecting reworking by tidal currents. O’Hara and Oldale (1987) have proposed that these post-

Late Quaternary construction of Cape Cod, Massachusetts

43

Figure 37. Topography (dashed contours) and isopachs (heavy contours) of post-transgressive Holocene sediments of the Monomoy Island region. Bathymetry from National Oceanographic and Atmospheric Administration chart 13237 and isopach contours from O’Hara and Oldale (1987, Fig. 8).

transgressive Holocene sediments were likely derived from the adjacent glacial drift. Their southwesterly trend is clear evidence that both Monomoy Island and its southwest extension, Handkerchief Shoal, were constructed from sediment derived from the north and transported to their present position by southerly littoral drift. If this is correct then only 0.28 km3 of the sediments in the Monomoy region came from the cliffs to the north and the rest of the detritus, 0.73 km3, used in the construction of Great Round, Stonehorse, and Little Round Shoals, was derived locally from erosion of the Pleistocene sediment during the postglacial Holocene transgression. The shoals have acquired their present east-west configuration subsequent to their creation as a result of tidal current action. Zeigler et al. (1965), who believed part of the detritus used to maintain the southern spits came in part from erosion of the nearby sea floor, estimated that this annual production was 184,974 km3, which converts to 1.1 km3 for the last 6,000 yr. Our calculations suggest, however, that this production may be much smaller and is

not essential for the maintenance of the spits as the cliffs to the north are capable of providing nearly 74% of the material needed (0.21 km3 of the 0.28 km3) to maintain the southerly spits. If our estimates are correct, then 86% of the material eroded from the cliffs is used in the construction of the Provincetown Hook, 7.3% in the constructions of the beaches and bars fronting the cliffs, and 6.7% in building the spits south of the cliffs. By contrast, Zeigler et al. (1965) postulated that roughly twice as much sand moved north along the lower Cape as moved south (458,760 m3 vs. 229,380 m3). An age of 6,000 yr for Provincetown Hook and other coastal constructional features in the lower Cape is not unreasonable. The oldest material closest to the base of the hook (9.1 m above the platform) is 4,375 ± 400 B.P. and estimates from the sand budget of the region indicate that the total loose sand in the lower Cape accumulated during the past 5,321 yr (Zeigler et al., 1965). The Provincetown Hook and the barrier islands at the elbow of Cape Cod were constructed from sedi-

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ments derived from the sea cliffs and the nearshore bottom and transported to their present depocenters by north- and southflowing littoral drift during the last 6,000 yr. Fisher (1979b, 1987) undertook a short-term synoptic collection of beach samples to attempt to locate the shoreline nodal point where the littoral drift diverged to the north and south, and the fulcrum point where erosion of the cliffs gives way to deposition on Provincetown Hook. A previous study by Schalk (1938), although collected at both high and low tide, was spread over two seasons, and the detailed study by Zimmerman (1963) was restricted to the region from Coast Guard Beach to Nauset Inlet. A similar study was undertaken during the present investigation where samples were collected both along the foreshore and backshore. Fisher took his samples at 300-m intervals, Schalk at an interval of about 1 km, and our samples are at a spacing ranging from about 3 to 5 km (Fig. 38). All three studies show that the beach sands in front of the cliffs display little textural variation but become coarser to the north, with the grain size increasing from 0.30 to 0.40 mm at the southern end to 0.55 to 0.90 mm at the northern end of the cliffs. Beach sands in the barrier spits south of the cliffs to Monomoy Point are quite variable in texture ranging in grain size from 0.65 to less than 0.3 mm. The sands in Provincetown Hook are the coarsest and display a uniform increase in median diameter from 0.55 to 0.90 mm at the eastern end to 1.5 to 2.3 mm at Long Point. Samples collected by

Schalk (1938) at low tide from Long Point to Pamet River are coarser than those collected at high tide. Only the low tide sand collected at Wood End Light depart from this trend being slightly finer than the high tide sample at the same locality. Samples collected during the present investigation also show that the foreshore sands tend to be slightly coarser than those from the backshore environment. Such textural variation is the result of the energy level from the swash to the surf zone with the highest being in the transition zone in the collision with the outgoing swash and the incoming surf (Komar, 1976, p. 351). Schalk (1938) reported that the finest grain size occurred along the cliffs and that the median diameter increased both to the north and south of the cliffs, a trend that Fisher (1987) suggested may be due to the fact that the samples were collected at different times. The samples collected by Fisher show that the sands increase from medium to coarse northward along the length of the cliffs, and that the sands in Provincetown Hook increase from coarse to very coarse along the spit. The overall trend along the east coast of the lower Cape is one of increasing grain size from about Nauset Beach to Long Point (Fig. 38). Thus, the grain size of the samples collected by Fisher north of Nauset after the summer season of prevailing southerly winds coarsen in the direction of sediment transport. The beach sands from Nauset south display an irregular textural pattern downdrift with no obvious change in grain size. That coarsening or fining in the direction of net transport does not occur south of

Figure 38. Texture of beach sands along the lower Cape. Based on analyses by Schalk (1938, Fig. 2), Fisher (1987, Figs. 6, 7), and the present investigation. The median diameters of the drift (D) in the cliff, the beaches (B), nearshore bar (NB), and offshore bar (OF) also are indicated.

Late Quaternary construction of Cape Cod, Massachusetts Nauset may be the result of local contributions to the sediment regime of the southern spits. Komar (1976, p. 355) listed four ways by which long-shore variations in grain size could be produced: (1) long-shore changes in wave energy; (2) selective rates of transport, where the fines are transported faster than the coarser grades; (3) winnowing of the fines from the beach, which are carried onshore by winds or offshore by waves; and (4) the interplay of waves from different directions. A fifth, and obvious, possibility exists: long-shore variation in source grain size. Although grain size does vary in the cliffs, it does not do so in a manner coherent with beach sand textural variations. McCave (1978) ascribed the increase in grain size in the direction of net transport in northern East Anglia in eastern England to the winnowing out of the finer fraction and its dispersal offshore by tidal currents, leaving a progressive coarser lag on the beach downdrift. He also suggested that in regions lacking strong tidal currents these fines could be reintroduced to the beach downdrift, leading to a fining trend downcurrent. Coarsening of the sands downcurrent along the lower Cape may be the result of a variation of the third mechanism proposed by Komar (1976) and the mechanism proposed by McCave (1978) for East Anglia. Grain-size analyses of the beaches and bars fronting the cliffs, the sediment source area, indicate that the coarsest sands, including gravel, accumulate in the beach and nearshore bar with the finer grades winnowed out by wave action to be deposited in the offshore bar. The beach and nearshore bar sands and gravel are subsequently transported laterally by the wave-driven longshore drift, eventually coming to rest in the distal end of Provincetown Hook. The finer, offshore sands are carried laterally to the south by net southward-directed longshore currents during stormy periods when they are put into suspension by breaking waves. Three models have been proposed for construction of Provincetown Hook. Davis (1896) proposed that the first sand spit began to build northwest of the glacial headland 4,000 yr ago. As the spits grew northwestward, they tended to curve and fuse at their ends, cutting off areas of water (Fig. 39A). Provincetown Hook built in this manner would tend to be quite narrow at the point where it is attached to the glacial highland and subject to inlet breaching as is Sandy Hook in New Jersey. In the Davis model the fulcrum of the spit shifted position as the glacial cliffs retreated southeastward. This shift was ascribed by him to the construction of a new tangential spit northward and seaward of the previous one. As described by Leatherman (1987), the Davis model is based on the premise that (1) the present-day curving dune ridges delineate the shoreline at different stages in the development of the hook and that the swales between the ridges represent areas of salt water that were cut off by the prograding sand spits; (2) the hook was built from east to west over the length of the Provincetown lowland before it was cut by the next spit being formed to its north; (3) the length of the hook was established with the formation of the first spit; and (4) subsequent spits constructed to the north

45

Figure 39. Evolution of the Provincetown Hook according to Davis (1896) (A) and Zeigler et al. (1965) (B). Adapted from Leatherman (1987, Figs. 4, 6).

increased only the width of the Provincetown Hook. In the modified Davis model proposed by Messinger (1958), the peninsula formed around a preexisting island or landmass off the tip of Cape Cod. Leatherman (1987) rejected the Messinger model, pointing out that drilling by Zeigler et al. (1965) had failed to yield evidence of such a high, and that the oscillations in sea level needed to build the beach-dune ridges in this model were not supported by the sea-level curve of Redfield and Rubin (1962). In a third model proposed by Zeigler et al. (1965), which is based on boreholes drilled in the hook and 14C dating of strata encountered during the drilling, the spits making up the Provincetown terrain were built sequentially, with each one being built a short distance to the west and then hooking abruptly to the south (Fig. 39B). In the model of Zeigler et al. (1965), the width of the hook also was established early in its history and its length increased slowly westward with time as a

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new spit was added to the feature. Zeigler et al. (1965) also believed that as a result of a slowing down in the rate of sealevel rise 2,000 yr ago, construction shifted from a pattern of tight hooks to deposition along the length of Provincetown Hook. Except for the recent sand ridges enclosing Hatches Harbor, the dune ridges also bear no relationship to the locations of the older shorelines in the model of Zeigler et al. (1965). To test the validity of the two models, Leatherman (1979f, 1987) took 41 cores in four swales in the Provincetown Hook to determine the stratigraphy of the lows. If the cores encountered plant remains and sediments of marine origin in the subsurface, which give way up section to a continental sequence as the area becomes more continental in origin, it would support Davis’s contention that the dune-ridge system delineates the location of a former shoreline. Similarly, the sediments on the swales on either side of a dune-ridge should have different ages for this concept to be correct. The core samples do not appear to support the Davis model. For example, no salt-marsh deposits were penetrated by any of the cores that yielded only a thin surface organic layer resting on an eolian deposit. Only at Race Run were marine sediments encountered beneath cranberry bogs. At three coring sites compacted peat deposits were encountered beneath eolian sands, which Leatherman (1979f, 1987) believed represented a mixed-shrub swamp similar to swamps found today in many areas of the Provincetown Hook. He further stated that, although the coring is not conclusive (only the upper 3 m of the 60-m section of Provincetown Hook were sampled), the orientation and characteristics of the relict freshwater marshes can best be explained by the model of Zeigler et al. (1965) rather than by the Davis model. The location of similar peat deposits in different swales rather than the same one, and the occurrence of marine coarse sand and pebbles 1 m below the surface in what according to Davis was supposed to be an old swale, also tend to question the validity of his model. All in all, the stratigraphy displayed by the cores taken in this complex geologic region of constant shifting dunes (even the stabilized dunes have been active in the recent past) is more compatible with the model of Zeigler et al. (1965). The various stages of retreat of the sea cliffs along the east side of the lower Cape and the development of Provincetown Hook are displayed in Figure 40, a history of continuous retreat of seaward-facing cliffs and the formation of sediment accumulations north and south of the retreating cliffs. About 9,500 yr ago, when sea level was 30 m lower, the lower Cape was dominated by an eastward-facing cliff extending from the High Head sea cliffs to Nauset (Fig. 40A). At the northern and southern ends of the cliffs at the present locations of Provincetown Hook and Nauset Marsh were two marine embayments. As the cliffs retreated westward and sea level rose to about –10 m 6,000 yr ago, the embayment south of the cliffs was filled with detritus eroded from the cliffs to the north and from material derived from the outwash plains surrounding the embayment. As the sediments prograded eastward, the southern tip of the cliffs in Nauset Marsh was slowly isolated from further marine

erosion. Sediment transport to the north then was nonexistent because the crest of Georges Bank rose above sea level up to 6,000 yr ago, inhibiting littoral drift in that direction. As the southern embayment became filled with sediment, the shoreline prograded eastward to or near the position of the shoreline along the seaward edge of the wave-cut terrace north of the embayment. We speculate that a barrier or system of barriers of unknown geometry was formed along the eastern edge of this filled low, constructional features that retreated westward, keeping pace with the rising sea (Fig. 40B). With the development of these barriers and marshes on their landward sides the cliffs defining the glacial terrains were protected from further erosion. By 6,000 yr ago, as Georges Bank drowned, sediment transport to the north began and construction of Provincetown Hook was initiated. Its subsequent buildup from 5,000 yr ago to the present is displayed in Figure 40, C–F. As Provincetown Hook grew, the cliffs at High Head, like those inboard of the spits at the Cape’s elbow, became protected from further marine erosion. During the last 2,000 yr Provincetown Hook has grown 1,800 m to the northeast at Race Point and some 600 m to the east of Pilgrim Lake (Zeigler et al., 1965), and in the last 70 yr, coastal accretion at the Race Point–Pilgrim Lake area has averaged 0.7 m a–1 (Zeigler et al., 1965). Geodynamics and historical changes. Although large expanses of the western Gulf of Maine east of Cape Cod Bay and the east coast shelf south of Martha’s Vineyard and Nantucket Island have not changed appreciably since they were inundated during the Holocene transgression, other segments of the Cape Cod region are in state of continuous metamorphosis (Fig. 41). Some of these transformations are slow, occurring within a framework measured in centuries to a millennium; others occur daily and are related to the tidal regime of the regions; and finally, features form at time intervals of less than a day to decades (see Aubrey and Giese, 1993, and references therein). Tidal current–generated terrains are short-lived geologically as they are in a continuous state of change in response to daily tidal cycles. Examples of such terrains are ebb-tidal and flood-tidal deltas and midestuary sand bodies as occur in the Chatham Harbor estuary (Hine, 1972), the features off the tip of Cape Cod and Stellwagen Bank, Nantucket Shoals, the shoals in Nantucket Sound, and Middle Ground in Vineyard Sound. The ebb-tidal delta, a multifacies asymmetrical body seaward of a tidal inlet, consists of sediment transported to the site seaward of tidal inlets; the flood-tidal delta landward of inlets is composed of flood spillover and complexes; and midestuary sand bodies are constructed of reworked flood-tidal deltas sequences (Fig. 41, B–C). Areas of strong gradients in tidal currents on Stellwagen Bank and around the tip of Cape Cod where the bottom sediments are intensively reworked by tidal currents also are areas where Right Whales are generally encountered (Blumberg et al., 1993; Hamilton and Mayo, 1988). Middle Ground, a sand ridge in Vineyard Sound, is a spit formed from sand eroded from Lucas Shoal during the Holocene transgression (Fig. 42A) (Smith, 1969). A six-day

Late Quaternary construction of Cape Cod, Massachusetts

47

Figure 40. Paleogeographic maps of the lower Cape showing the successive positions of the shoreline 9,500 (A), 6,000 (B), 5,000 (C), 3,700 (D), 2,000 (E) B.P., and today (F). Compiled from Zeigler et al. (1965, Figs. 5, 6, 8) and other sources discussed in the text. Sea level is relative to the present level.

series of photographs taken with a time-lapse camera in Middle Ground quite clearly shows movement of the sand on the bottom due to the alternating currents (Owen et al., 1967). Some of this sand leaves the system at the northern end of the shoal at West Chop, but much of it is recirculated by the tidal currents down the northern side of the high. This counterclockwise sediment transport around Middle Ground has been supported by

more recent work by Briggs (1979). The most extensive of these tidal current–generated terrains is Nantucket Shoals, dominated by its shallow northeast-trending sand ridges aligned parallel to the tidal currents. During the flood and ebb tides the crests of the sand ridges are reworked by the currents and the finer fraction is continuously winnowed from the highs. Once suspended, they become entrained in the water column (a

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Figure 41. Morphology of the recurved spit complex (A), ebb-tidal delta (B), and flood-tidal deltaebb spillover lobe complex in the inlet between Monomoy Island and Nauset Spit (C). Compiled from Hine (1972, Figs. 2, 24, 49, 51). Maps are at different scales so as not to lose detail.

mixture of Gulf of Maine and Nantucket Sound waters) that leaves the region via Great South Channel. Here this water and its suspended load mixes with shelfwater and joins the general westward flow over the mid- and outer east coast shelf (Limeburner and Beardsley, 1982). During this westward transport the fines winnowed from the Nantucket Shoals ridges tend to

deposit in the mud patch south of Martha’s Vineyard. Some of the material winnowed from Nantucket Shoals never leaves the region of the shoals but spills over their eastern edge to accumulate on Fan A. Features formed at intervals of hours to years are usually constructed by littoral drift generated by waves or residual tidal

Late Quaternary construction of Cape Cod, Massachusetts

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Figure 42. A, Chart showing locations of Lucas Shoal and Middle Ground, a sand spit. B, Changes in the position of the crest of Middle Ground Shoal with time. Positions of the crest of the shoal are based on National Oceanographic and Atmospheric Administration smooth sheets and aerial photographs. Compiled from Smith (1969, Figs. 2, 8).

currents. Some of these terrains form slowly, with their development best documented by surveys repeated over periods of years. Other features, however, develop in a matter of hours as a result of some catastrophic event, generally an intense storm. Some of the wave-generated terrains are short-lived, having been formed during a tidal cycle that will be modified to some degree during the next tidal cycle (Fig. 43). Another example of a more lasting type of terrain is the migration of the shoals due to residual tidal currents. A comparison of topographic surveys made in 1886, 1906, and 1962 shows that the crest of Middle Ground has bent with time (Fig. 42B). Surveys of Middle Ground made by Briggs (1979) in March, June, August, and November 1978 also indicated not only that the whole shoal has shifted, but that the sand waves in the ebb-dominated southern flank migrated to the northeast. Other sand waves, instead of migrating, tend to flex their crests like the shoal itself, flexures that reverse themselves during the tidal cycle. Some of these changes appear to be seasonal and are the result of wave action that is more intense in the winter months. According to Briggs (1979), the yearly average rates of the migrations caused by residual tidal currents ranged from nearly zero to about 40 m a–1. In Nantucket Sound two of the sand shoals themselves may have migrated from their original position and come to rest over a marsh deposit and Holocene marine sediments more than 10,000 yr old. Other examples of long-term change in terrain morphology

can be seen in the lower Cape. Comparison of topographic maps of Wellfleet Harbor from 1887, 1933, and 1958, for example, show the disappearance of Billingsgate Island, which submerged in 1942, and the evolution of the spit along the seaward side of the harbor, an evolution marked by erosion when the spit was cut in two, and subsequent growth to its present shape (Fig. 44). Destruction of the island was in part due to dwindling of southerly transport as a result of the development of Provincetown Hook. Similar surveys of Pamet River show the artificial cut made in the spit and its northern growth with time (Fig. 45). In the early 19th century, Pamet Harbor was a shipbuilding and fishing port. According to Fisher (1979a), a lighthouse was erected at the entrance of the harbor in 1879 but was discontinued 6 yr later when the harbor silted up. Pamet River is separated from the open ocean at its eastern end by the Ballston Beach–dune complex (Fig. 46). This beach-dune system has been overwashed several times during severe northeast storms (1972, 1978, and 1992). Analyses of historical maps also show the changes that have taken place along the western side of Pilgrim Lake, formerly East Harbor (Fig. 47). Comparison of surveys made in 1833–1835 with those of 1867 shows the changes in the configuration of the spit along the seaward side of the harbor. This barrier, formed by littoral drift to the northeast, was enhanced with the development of Provincetown Hook, and was later artificially extended northward in 1869 to connect it with the

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Figure 43. Wave-generated sand bodies in Cape Cod Bay exposed during low tide. The view is looking south with the Boat Meadow River in the foreground and the entrance to Rock Harbor in the background. These features are transitory and will change shape during the next high tide.

hook (Mayflower Heights) and stabilized with the planting of beach grass. The connection later was further strengthened with fill to provide a causeway, first for a railroad to Provincetown and more recently for two highways (Leatherman, 1979b; Giese and Leatherman, 1979). With its closure, the former harbor, now a lake, became progressively less salty with time. Similar changes also have been noted in the periphery of Provincetown Hook, with its eastern side undergoing retreat and its western side accreting (Fig. 48). The entrance to Hatches Harbor at the western tip of Provincetown Hook also has undergone noticeable changes during the last century. During this time the harbor has been almost closed by the growth of the recurring spits, and only a narrow inlet has survived to provide a passageway to the open ocean (Fig. 48) (Hamilton, 1978; Leatherman, 1979c; Hamilton and Leatherman, 1979). Hamilton (1978) reported that from 1841 to 1887 the harbor was 4.4 km long and 1.2 km wide, with numerous islands rising above high tide, and was probably about 2.8 m deep. By 1910, infilling had reduced the depth of the harbor to less than 1 m, but its inlet deepened and broadened with time, reaching a width of 0.36 km in 1944. Since then the inlet has again decreased in width to 0.1 km, migrated southward, but at the same time maintained its offset shape. In time, this low will be infilled with sediment to become one of the interdunal lows characteristic of the Provincetown Hook, with a small tidal inlet connecting the marshy area with the open sea. Changes in dune and bog and pond morphology in Provincetown Hook appear to be the result of climatic changes and human activity. Modern wetlands were initiated about

Figure 44. Shoreline maps of Wellfleet Harbor showing historical changes of the spit and Billingsgate Island along the west side of the Harbor. Based on charts from U.S. Coast and Geodetic Survey.

Late Quaternary construction of Cape Cod, Massachusetts

Figure 45. Shoreline maps of the Pamet River showing historical changes at the western end of the Pamet River. Artificial refers to manmade features such as dikes, culverts, and fill. Based on charts from U.S. Coast and Geodetic Survey.

Figure 46. Panoramic view of the eastern end of Pamet River Valley. Ballston Beach separating the valley from the open ocean has been overwashed a number of times during storms, with the gap produced providing a temporary connection between the valley and the sea. These photographs were taken in December 1992 by G. S. Giese, Woods Hole Oceanographic Institution, after the latest overwash that occurred in October 1992, documenting that within several months after it was eroded, the breach had already been healed.

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Figure 47. U.S. Coast and Geodetic Survey maps of East Harbor (present Pilgrim Lake) showing historical changes of the region. The barrier separating the lake from the open ocean is artificial fill and was constructed in 1869.

1,000 yr ago in the central and southwestern part of the hook within an earlier dune field. Winkler (1992) suggested this wet episode was due to a combination of factors, which include a slowing down of sea-level rise, a slight change in the direction of the prevailing wind, an increase in precipitation that led to an increase in the vegetation cover and subsequent stabilization of the dunes, and/or a slowing down of sand-laden winds. Peatlands also developed within the present active dune field from 700 to 500 yr ago as a result of increased precipitation and decreased wind activity. As a result of more intense, colder and drier winds, dune activity was renewed during the Little Ice Age from 500 to 100 yr ago, an activity that was accentuated by the arrival of European settlers to the region who denuded the dunes for fuel and agricultural purposes. The eolian activity in

the region again diminished during the last century as a result of a change to a wetter, warmer, less windy climate that favored the development of wetlands and the stabilization of the dunes by revegetation through human efforts. Modifications on the western side of the lower Cape are not limited to changes in the configuration of Holocene constructional features, however, but also to changes in the configuration of the glacial terrain. Like the cliffs on the glacial deposits on the Atlantic side, those on the Cape Cod Bay side also are retreating, but at a much slower rate of 0.3 m a–1 (Giese, 1964; Leatherman, 1979a). From August 1887 through 1889, Marindin (1889, 1891) surveyed 229 points along most of the east coast of the lower Cape including Provincetown Hook. Reoccupation of 74 of those points in 1958 and 1959 (Zeigler et al., 1964b) demon-

Late Quaternary construction of Cape Cod, Massachusetts

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Figure 48. Historical changes in Hatches Harbor. From Hamilton (1978, Figs. 5 through 7).

strated that the Provincetown Hook from Race Point to Pilgrim Lake is accreting at a rate of 0.70 m a–1, and that the cliffs from south of Pilgrim Lake to Nauset are retreating at a rate of 0.8 m a–1. Comparison of aerial photographs taken in 1938, 1952, 1971, and 1974 (Fig. 49) of the coast of the lower Cape from Long Point to Monomoy Point indicates that the greatest changes of the high water line were at the northern and southern ends of the coast (Gatto, 1979). At the north end net changes varied from –3 m (erosion) to +97.9 m (accretion) and at the southern end from –18.9 to –547.3 m; net changes in the central segment ranged from –2.1 to –48.2 m. These and similar historical records suggest that the cliff is retreating at a rate of about 0.8 m a–1; this, together with the supposition that sea level has been near its present level since about 3,500 yr ago, have been used to determine the shape of glacial lower Cape. Fisher (1972) estimated that the shoreline extended 3.7 km offshore, Shaler (1897) calculated that it was no less than about 1 km and no more than 7.2 km, and Davis (1896) thought that it was about 4 km. Our estimates, based on seismic reflection data seaward of the lower Cape, indicate that the glacial lower Cape extended 2.8 to 7.0 km east of the present shore. Cliff retreat from this position to its present one probably was not at a uniform rate, but was marked by periodic catastrophic failures

such as the one that occurred in 1848 when more than 18 m of land disappeared from in front of the Highland Light, and in 1978 when another 5.5 m broke off (Fisher and Leatherman, 1979). Assuming an average rate of cliff retreat of 0.8 m a–1, the lower Cape—which ranges in width from 2.4 km in the north, 7.0 km in the center, and 5.0 km in the south—should be eroded by 3,000 years at its north end, 8,800 yr in its center, and 6,200 yr at its southern end. Once the lower Cape is destroyed, the upper Cape will soon follow and the former position of the Cape and environs will be marked by an extensive sand plain and shoals and scattered patches of lag gravel. The elbow of Cape Cod with its spits and extensive marshes also has undergone noticeable changes in historical times. Nauset Beach, which dominates the region, for example, is slowly being displaced landward by tidal dynamics and overwash processes (Leatherman, 1979d, 1979e). A map compiled by Captain Cypian Southhack in 1717 showing the location where the 30.5-m brigantine Whidah sank April 26, 1717, displays a navigable channel at the Cape’s elbow connecting Cape Cod Bay and the Atlantic, an area occupied by marshes today. This path probably represents remnants of the gap along which Cape Cod Bay Glacial Lake drained into the Atlantic toward the end of the Wisconsin. Although the passage may have been

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Figure 49. A, Net annual rate of change in the high water line on the lower Cape from Long Point to Monomoy Point. B, Locations of reference points in A. Modified from Gatto (1979, Figs. 1, 4).

navigable at high water, the present distribution of the glacial deposits in the region indicates that the waterway was much narrower and more irregular than indicated on the Southhack map. Changes in the morphology of Nauset Inlet breaching Nauset Spit are documented by historical charts as old as 1605. A chart drawn by Champlain in 1605 (Ganong, 1922) shows the southern spit at that time as being located 0.8 km seaward of its 1978 location, suggesting that from 1605 to 1948 the spit retreated at a rate of 2.1 m a–1 (Wright and Brenninkmeyer, 1979). Analyses of historical charts and photographs by Aubrey and Speer (1984) indicate that Nauset Inlet has migrated extensively during the last 30 yr with its preferred location being just north of Nauset Heights. Prior to 1944, the south spit does not appear to have been significant, and from the 1950s to early 1980s the inlet displays three cycles of northerly drift with the

north spit being stable and the southern one growing northward (Fig. 50). Many of the changes documented by the charts were storm induced when the northern spit was breached to form a new inlet. Studies by Wright and Brenninkmeyer (1979) from June 1975 to June 1976 indicate that after the northern spit is breached during a storm the spit remnant south of the breach develops into a flood-tidal delta. According to Aubrey and Speer (1984), three mechanisms caused this updrift migration of the tidal inlet: (1) attachment of distal ebb tide delta bars to the downdrift spit; (2) storm-produced breaching of the spit followed by stabilization; and (3) channel discharge producing a three-dimensional flow regime leading to erosion on the outer channel bank and deposition on the inner channel bank (as in stream meanders). Analyses of historical charts show similar changes in Nauset

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Figure 50. Historical charts showing the migration of the Nauset Inlet with time. Compiled from Wright and Brenninkmeyer (1979, Fig. 1) and Aubrey and Speer (1984, Figs. 4, 6, 7). In the 1605 chart, the ∗ marks the location of Nauset Heights and the letter H is a segment of upland terrain that Ganong (1922) interpreted as a small drumlin.

Beach in front of the Chatham Lighthouse and Monomoy Island (Fig. 51). Nauset Beach and Monomoy Island are two barriers with tidal flow between Chatham Harbor–Pleasant Bay Estuary taking place through South Inlet located south of Nauset Beach. Comparison of maps from 1887, 1940, 1947–1953, 1961–1964, with stereoscopic photographs taken in 1969 indicated that the northern end of Monomoy Island (Shooters Island) has been receding since 1948 and that it could have separated from the rest of the island in about 70 or 80 yr (Oldale et al., 1971). The break, however, took place in 1978, earlier than anticipated by Oldale et al. (1971). The southern half of Monomoy has been prograding eastward at a rate of about 12 m a–1 since at least 1853. Growth of Monomoy will probably slow down in the future as the island progrades into Butler Hole to the southeast and a smaller depression southwest of the island. As Nauset Beach propagates southward, Monomoy Island tends to separate from Morris Island, leading to the creation of a new tidal inlet, West Inlet, north of Monomoy. As

Nauset Beach continues to grow southward, a new inlet is breached in the spit with the segment south of the breach migrating onshore to fill in the old tidal channel. In time, this migrating sand seals the opening north of Monomoy connecting the island to the mainland again (Friedrichs et al., 1993). Such a breaching of Nauset Beach occurred during an intense northeast storm on January 2, 1987 (Fig. 52). As a result, houses immediately to the west of the ever-widening breach, which are no longer protected by the spit, have incurred and continue to incur intense damage during storms. Inlet creations, inlet displacements, westward retreat of the barrier islands, and southerly spit extensions at the Cape’s elbow are documented by charts going as far back as the early 17th century. McClennen (1979) believed that these changes are cyclical, that they occur every 100 yr and that cyclicity corresponds to a period when water levels in Pleasant Bay differ from those in the Atlantic by 1 m. More recently, Friedrichs et al. (1993) constructed a branched one-dimensional numerical

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Figure 51. Historical changes in the Nauset Beach–Monomoy Island barrier system. Compiled from Giese (1988, Fig. 2) and Friedrichs et al. (1993, Fig. 2). SP = Strong Point; CBI = Chatham Bars Inlet.

model to attempt to determine why the barriers and estuaries exhibit a long-term cycle of inlet formation that Giese (1988) postulated to be about 150 yr, and why once an inlet is formed the resulting multiple inlet system is unstable. On the assumption that these episodic changes in the spit will continue into the future, Giese made some predictions for the periods 1975–1985, 1985–1995, and 1995–2005 for the Chatham Conservation Commission in 1978. He predicted that a new breach would form on the spit sometime during the 1985–1995 decade and that the breach would occur east of North Chatham; the break, which occurred in 1987, took place 3.9 km south of his predicted location. Friedrichs et al. (1993) also proposed that a new inlet would be formed in Nauset Beach when the elongation of the spit led to the creation of a critical hydraulic head across the barrier as previously suggested by McClennen (1979). If the critical head, enhanced by storm surges and wave set-up, occurred at the time of high tide during consecutive tidal cycles, then the storm currents and waves would tend to deepen the overwash channel, allowing the stronger ebb currents to complete the formation of the inlet. With the creation of this new inlet, tidal discharge through the older one will be diminished and it will become a site of deposition. As the region of the lower inlet shoals, the northern and southern sections of the Chatham Estuary will become hydrodynamically decoupled.

As the spit begins to grow southward again, ultimately it will reach a critical length, building the required hydraulic head in the estuary, and will initiate a new cycle of spit breaching. An example at the slow end of the geologic time scale where changes in the Cape region are imperceptible over short time periods and occur in centuries to millennia is Cape Cod Bay (Fig. 53). This body of water within the bent arm of Cape Cod has acted as a sediment sink for sediments from the north throughout much of the Holocene. Neither Cape Cod Bay nor Massachusetts Bay to the north has a sizable river discharging directly into it, but discharge from the Merrimack River and other Gulf of Maine rivers north of Cape Ann generate a buoyancy-driven southerly flowing coastal current that transports sediments from those rivers to Cape Cod Bay. This flow is part of a counterclockwise circulation in the region made up of southerly inflow south of Cape Ann, southerly flow off Scituate, and northeasterly flow north of Race Point (Bigelow, 1927; Bumpus and Lauzier, 1965; Brooks, 1985; Blumberg et al., 1993). According to Geyer (1992), this residual flow pattern reverses in the fall. Most of the material in suspension in these coastal waters tends to bypass Massachusetts Bay and accumulates in Cape Cod Bay accounting for its smooth topography relative to the rougher terrain in much of Massachusetts Bay (Fig. 54). Some of the sediment does leave Cape Cod Bay north

Late Quaternary construction of Cape Cod, Massachusetts A 1980

B 1991

Figure 52. Photographs of the Chatham Break that took place on January 2, 1987. Photograph A is from the site taken in 1980 prior to the break. Photograph B was taken in the summer of 1991 at low tide when the breach was nearly 2 km wide. Note flood-tidal delta blocking the former south tidal inlet. Photograph A courtesy of A. G. Gaines, Woods Hole Oceanographic Institution; photograph B courtesy of D. S. Blackwood, U.S. Geological Survey, Woods Hole.

of Race Point. Once the sediment escapes the confines of the coastal bays, it is transported southward by the Gulf of Maine water; as it is transported southward, the sediments accumulate in the western Gulf of Maine and beyond. Historical changes in Cape Cod are not limited to natural events but are the result of human activities. On a more global scale, Hooke (1994) reviewed some of the impacts of humans as geomorphic agents. Many of these impacts occur on Cape Cod. Included among these activities are devegetation of the dunes of Provincetown Hook by off-road vehicles (Leatherman, 1979g) and mining sand for more than 50 yr in the Provincetown Hook (Leatherman and Godfrey, 1979), practices no longer allowed. Timber harvesting, clearing of land for agriculture, and unrestricted grazing from the colonial period to the

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recent past are other practices that have severely altered Cape Cod’s ecosystem (McCaffrey and Leatherman, 1979). This misuse of the land led to an increase in eolian activity and the shoaling of Provincetown Harbor by the deposition of sand. Such deposition in East Harbor in the middle of the 19th century made it too shallow for continued use as a harbor of refuge (Leatherman, 1988, p. 96). Removal of the vegetation cover may also have led to the erosion of a channel during Minot’s Gale in 1851 connecting East Harbor to the Atlantic via the Salt Meadow at the foot of the High Head cliffs (Leatherman, 1988, p. 96–97). This breach, which some thought threatened to destroy Provincetown Harbor, led to the closing of East Harbor and the creation of Pilgrim Lake and the planting of beach grass to heal the breach. Building of the dike has led to other problems, including the formation of a highly eutrophic lake unsuitable for recreation. Leakage from the tidal gate controlling the outflow of the water from the lake led to saltwater influx and a fish kill in 1972. This in turn led to a massive increase in the midge fly population, which threatened the tourist industry. To prevent similar problems in the future, the tidal gate was improved (Leatherman, 1988, p. 97). Another human-made feature that has affected the ecosystem of Cape Cod is the Cape Cod Canal, 32 km long, 146 to 231 m wide, and about 10 m deep. This artificial feature, completed in 1936, affected the littoral drift in Cape Cod Bay (accelerating erosion south of its eastern entrance) and led to an eco-linkage between Buzzards Bay and Cape Cod Bay. The west side of Cape Cod (Buzzards Bay) has been modified by a variety of human activities as well as natural erosive processes. Among the major changes to the west side of Cape Cod are the construction of numerous harbors along the coast (Phinney’s Harbor, Pocasset Harbor, Red Brook Harbor, Squeteague Harbor, Megansett Harbor, West Falmouth Harbor, and Quissett Harbor: Fig. 2). These harbors are located in reentrants to the morainal topography of western Cape Cod, altering freshwater exchange with the bay. In one instance (West Falmouth Harbor), a natural entrance was closed to maintain a strong tidal flow in the remaining entrance. The stronger flow was desired to maintain a self-sustaining navigation channel to permit passage of pleasure craft. The other changes include dredging of entrance channels and mooring areas, and construction of harbor facilities along the shoreline. Although beach erosion along this section of Cape Cod is less than the more exposed Atlantic shore of Cape Cod, it is still significant, particularly along the many headlands. Protection of these headlands by hard engineering structures has reduced the sediment supply to adjacent beaches. An example is the armoring of Gunning Point and Hamlin Point by large rock revetments, cutting off sediment supply and turning Flume Pond into a retreating and thinning cobble beach instead of its more natural sandy texture. The south shore of Cape Cod has experienced significant changes in time, partly due to human influence and partly to

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natural forces. The tides in Nantucket Sound are relatively small (range of about 30 to 75 cm) compared to the Buzzards Bay shore (range of about 1 to 1.5 m), the Cape Cod Bay shore (about 3-m range), or the open Atlantic (nearly 2-m tidal range). Also, the waves are the smallest among the ocean shores of Cape Cod, all the waves being locally generated because of

complete encirclement by shallow water which precludes free passage of open ocean waves (e.g., Goud and Aubrey, 1985). This combination of small tidal range and low wave activity makes the south shore of Cape Cod relatively calm compared to the other Cape Cod shores. However, changes have taken place along this shore, particularly in the barrier beaches. One exam-

Figure 53. Holocene sedimentary domains of Cape Cod.

Late Quaternary construction of Cape Cod, Massachusetts

Figure 54. Computer-generated relief diagram of Cape Cod Bay and vicinity. Figure courtesy of R. P. Signell, U.S. Geological Survey, Woods Hole. Note smoothness of the sea floor of Cape Cod relative to Massachusetts Bay to the north. Lack of relief in Cape Cod Bay is the result of faster rates of sedimentation to the south. Note 100× vertical exaggeration.

ple is Popponesset barrier beach in Mashpee, which has experienced a cycle of elongation, shortening, and elongation in response to tidal inlet formation. Although much less dynamic than the Chatham situation, the Popponesset breach is a significant one for local water quality (Aubrey and Gaines, 1982). The historical evolution shows a lengthening barrier beach from about 1844 to 1954, when a series of hurricanes breached the elongated barrier about halfway along its length. This breach became the preferred inlet location, separating the northern island segment from the shore-attached southern segment. The north spit eventually migrated onshore, filling in the former inlet channel, disappearing completely as it fully filled the old channel. Another significant modification of the southern Cape Cod shoreline arose from navigation and harbor projects. Nearly two dozen significant harbors are scattered along the coast of southern Cape Cod. Most are protected by double jetties to minimize shoaling at the entrances and simplify boat passage. These jetties have altered the natural flow of sediment transport along the coast, causing accretion on updrift sides of the inlets

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and massive erosion on downdrift sides of the inlets. The satellite image of Cape Cod (Fig. 55) demonstrates these large inlet offsets, a dominant feature of the south shore of the Cape. The altered sediment budget has led to significant erosional problems, encouraging the use of revetments and seawalls to protect threatened structures. These hard engineering solutions have further reduced the sediment supply to the south Cape, such that this region is sediment starved at present. Although three beach nourishment programs have been conducted in south Cape Cod in recent years to help retard the erosion of the Cape, there is still insufficient erodible sand to reproduce the wide sandy beaches of this region in previous years. Probably the most pronounced impact on the fragile Cape has been the marked increase in the summer population, particularly during the last four decades. This summer population explosion, in turn, has led to a boom in various types of services from hotels, restaurants, and recreation facilities. There also has been a major increase in year-round population as more people see the Cape as an ideal place to live and retire, and with this has come an increase in housing and expansion in public services that such an increase entails (Fig. 56). As communications with Boston improve, this year-round population will certainly increase as more people leave the city to live in the Cape’s more hospitable environment. With them also will come many problems they believe they are leaving behind. With the increase in land use has come all its shortcomings: increase in waste products and the problem of what to do with them, contamination of the environment including the groundwater (the only source of drinking water for the Cape), attempts to engineer the environment to maximize its recreational potential, and the demand for instant gratification with little interest for the future consequences of such actions. It is the human condition to accept solutions that cause the least inconvenience, rather than accept those that may entail some sacrifice. Secondarily treated sewage disposed on surface sand beds since 1936 at the Otis Air Force Base (now called the Massachusetts Military Reservation) sewage treatment facility, for example, has infiltrated the sands into the underlying unconfined sand and gravel aquifer, producing a 23-m-thick, 760- to 1,970-m-wide, and >3-km-long south-southwest– trending plume of sewage-contaminated and otherwise polluted water (LeBlanc, 1982). The plume extends in the direction of the regional flow of groundwater, and is overlain by 6 to 15 m of groundwater produced by precipitation recharge of the aquifer. Although it was not sampled farther than 3.5 km from the treatment plant, the south- and southwest-moving plume is probably flowing in the direction of the downstream end of the Coonamessett River in East Falmouth and the small streams, ponds, wetlands, and saltwater bays east of the river 5 km southwest of the head of the plume (LeBlanc, 1982). Several other chemical plumes are also present within and outside the base boundaries. It is estimated that $250 million will be needed to restore the contaminated aquifer to potability. Those

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Late Quaternary construction of Cape Cod, Massachusetts Figure 55. Landsat image of Cape Cod showing the present extent of the late Pleistocene (late Wisconsin) glacial terrain, the terrains accreted onto this glacial landscape by coastal marine processes during the Holocene, and severe sea-ice conditions on February 15, 1979. Annotated image courtesy of U.S. Geological Survey, Quissett campus, Woods Hole (Williams et al., 1981).

of us living in the region must always keep in mind that the Cape possesses only finite natural resources and those who use it today have an obligation to leave it in a condition that future generations can also enjoy. This is particularly true of the aquifer that underlies the peninsula of Cape Cod, a sole-source aquifer that provides virtually all of the home and town water supplies from wells or ponds. CONCLUSION The raw data provided by geologic and geophysical studies of Cape Cod and southeast coastal Massachusetts have been used in this investigation to reconstruct the huge jigsaw puzzle that constitutes the geologic history of the region. This extensive panorama consists of three distinct events: fluvial, glacial, and marine. During the first event, a cuesta/lowland terrain was carved out of the shelf strata during marine regressions that go back as far as the Oligocene. The cuesta that extends from Long Island, New York, in the west to the Grand Banks of Newfoundland in the east is represented by Long Island, Block Island, the Elizabeth Islands, Martha’s Vineyard, Nantucket Island, Georges Bank, and the offshore banks on the Scotian Shelf off Nova Scotia and the Grand Banks of Newfoundland, Canada. The lowlands associated with this cuesta complex consist of Block Island and Nantucket Sounds, the Gulf of Maine,

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the basins on the Scotian Shelf, the Gulf of Saint Lawrence, and the shallow basins along the inner edge of the Grand Banks. In the Cape Cod region, coastal plain remnants reflecting these erosional events extend to the northern and western edges of Cape Cod Bay, beneath Stellwagen Bank and the western Gulf of Maine. Water gaps of the fluvial system that carved this fluvial terrain are represented by Vineyard Sound, Great South and Northeast channels on either side of Georges Bank, and the Laurentian Channel separating the Scotian Shelf and the Grand Banks. Although this fluvial terrain was apparently glaciated more than once during the Pleistocene, only the late Wisconsinan glaciation appears to have had any significant effect on the fluvial terrain. In the second event, late Wisconsinan South Channel, Cape Cod Bay, and Buzzards Bay ice lobes appear to have attained their maximum extent about 20,000 yr ago, reaching Martha’s Vineyard, Nantucket Island, and the northern edge of Georges Bank. The ice may have reached the open sea via Great South and Northeast channels on either side of Georges Bank. The ice began to retreat northward no later than 18,000 yr ago as marine conditions were established in the western Gulf of Maine no later than 17,600 yr ago, and by 14,000 yr ago the Cape Cod region was ice free. Prior to its retreat, subglacial streams carved a complex valley system out of the coastal plain sediments in Nantucket Sound. The hydraulic head in these waters was high enough that erosion extended up the back slope of the cuesta. Cape Cod construction, from river discharge from the retreating ice front, took place from about 18,000 yr to before 14,000 yr ago. Periodic readvances of the ice lobes led to the deformation of the sediments in front of them to form the moraines in Martha’s Vineyard and Nantucket Islands and in the upper Cape. We concur with Oldale and O’Hara (1984) that Billingsgate Shoal in Cape Cod Bay also

Figure 56. Population trends on Cape Cod from 1765 to 1990. Data provided by D. Finn, Cape Cod Commission.

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had such an origin and was not formed by littoral drift during the Holocene, as suggested by Davis (1896, 1954) and Giese (1963). The upper Cape sediment wedge was deposited from southerly discharge from the Cape Cod Bay lobe, a southerly dipping sediment wedge that graded to a lake that existed in Nantucket Sound at that time and reached a maximum thick-

ness of about 120 m near the southern coast of the upper Cape (Fig. 57). Westerly discharge from South Channel lobe deposited a sediment wedge extending from the northern tip of Stellwagen Bank to Great South Channel. At the southern end of the wedge, Nantucket Shoals rests on the coastal plain sediments. At the northern end of the depocenter, another local sediment

Figure 57. Isopach map, in meters, of the late Wisconsinan glacial deposits of Cape Cod and vicinity. Compiled from Oldale (1969), O’Hara and Oldale (1987, Fig. 5), Oldale and O’Hara (1980, Fig. 5; 1990, Fig. 7), Oldale et al. (1973, Plate 2B), and seismic reflection profiles recorded by Ballard and Uchupi (1975) in the western Gulf of Maine and archived at the Woods Hole Oceanographic Institution. The stippled pattern indicates areas where sediment thickness exceeds 80 m.

Late Quaternary construction of Cape Cod, Massachusetts accumulation forms the lower Cape. Sediments forming the lower Cape are graded to a lake that existed in Cape Cod Bay at that time and reached a maximum thickness of more than 160 m near the ice front and at High Head cliffs (Fig. 57). Partial collapse of these sediments led to the formation of a lake east of the lower Cape. The drifts damming the glacial lakes in Cape Cod Bay and east of the lower Cape were highly permeable, leading to seepages from the lakes behind them. Runoff from these seeps led to the erosion of the linear valleys in Cape Cod, valleys that migrated up the plains and terminated when the water table became too deep to be an effective erosive agent. Valleys in the upper Cape drained south toward Nantucket Sound and those in the lower Cape westward in the direction of Cape Cod Bay. Cape Cod Bay Glacial Lake drained first southwestward via Monument Valley, then southward via Bass River, and then southeastward via a gap between the clastic wedges forming the lower and upper Cape. How and in what direction the lake east of the lower Cape drained are yet to be determined. As the ice retreated northward, large expanses of New England that had been depressed by the weight of the ice sheet were flooded by the advancing sea, reaching elevations more than 30 m above the present level about 14,000 yr ago. As the regions rebounded isostatically about 12,000 yr ago when the weight of the ice was removed, the sea regressed and dropped to a level as much as 55 m below the present one. From there the sea level rose again, approaching its present level about 1,000 yr ago. Relative sea level is still continuing to rise at a rate of 2 to 3 mm a–1. Glacial Cape Cod had an area of about 1,562 km2 and extended eastward as much as 7 km from its present position. In contrast, Davis (1896) postulated that the glacial Cape extended only about 4 km eastward from its present position. During the third event, submergence by rising sea level and marine erosion have reduced the glacial Cape to an area of 1,077 km2, or 69.0% of its original size. Today the Cape is made up of glacial deposits covering an area of 1,077 km2 (88%) with the remaining 140 km2 (12%) consisting of the Barnstable marsh and spit, the Provincetown Hook and the bars, spits, and marshes. This composite origin of Cape Cod is clearly defined by the Landsat image displayed in Figure 55. The glacial terrain is scarred by numerous kettles, kettle lakes, dry valleys of spring sapping origin, and morainal ridges of glaciotectonic origin, giving the terrain a hummocky texture. Extending from this glacial core are coastal constructional recent marine features whose linearity clearly displays their origin. Aeolian forms capping both the glacial and marine terrains added further complications to the geomorphology of Cape Cod. The Holocene sediments rarely exceed 10 m in thickness and only in Provincetown Hook and east of Cape Cod are there appreciable accumulations of that age (Fig. 58). Davis believed that the reduction of the glacial Cape and the formation of the present one took place in about 4,000 yr, but our investigation indicates that this event lasted 9,500 yr. Davis

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also proposed that material eroded on the lower Cape was transported northward to form Provincetown Hook, but our investigations suggest that 82% of the 6.1 km3 (5.0 km3) sediment volume eroded from 9,500 to 6,000 yr ago (the remaining 18% was lost offshore), when sea level rose from –30 to –10 m, and was transported southeastward not northwestward. At the time Georges Bank was partially exposed and littoral drift in the direction of Provincetown Hook was inhibited. Sediments transported southeastward were deposited in a depression near the Cape’s elbow. This detritus provided about 64% of the depression’s fill with the remainder coming from the surrounding glacial plains. From 6,000 yr to 1,000 yr ago, after Georges Bank finally drowned about 6,000 yr ago, the direction of littoral drift shifted toward the northwest and sediments eroded from the sea cliffs in the lower Cape were transported in that direction to construct Provincetown Hook. Of the sediment eroded from the sea cliffs during the last 6,000 years and available for the construction of the nearshore features (3.28 km3), 2.8 km3 (86%) was transported northward to construct Provincetown Hook, 0.2 km3 (7.1%) to construct the beaches and offshore bars fronting the cliffs, and 0.2 km3 (6.9%) to construct the spits south of the cliffs. Today the cliffs along the eastern side of the lower Cape are retreating at a rate of 0.8 m a–1 (Zeigler et al., 1964b), a retreat that is enhanced not only by the erodibility of the material making up the cliffs, but also by a relative rise in sea level of about 2 to 3 mm a–1. If such a retreat were to continue for the next 20,000 yr, the Cape would be destroyed and its former position would be marked by extensive sand plains, sand shoals, and lag gravels. Part of the relative rise in sea level, about 1 mm a–1, is believed to be eustatic in origin, with the remainder being the result of some tectonic process. It has been suggested that this tectonic rise in sea level in formerly glaciated regions may be due to the collapse of the peripheral bulge that migrated northward with the retreating Wisconsin ice sheet. Such a tectonic origin for the noneustatic part of the rise in sea level in New England, a region that was deglaciated more than 12,000 yr ago, may be unrealistic. The Atlantic margin of North America, like other passive margins, is located along the transition of the first-order relief features of the Earth’s surface, the continental plateaus, and oceanic basins, a structural position of considerable instability. The Atlantic margin’s geologic history has been generally one of subsidence since its inception over 200 Ma, a regime that was complicated further by pronounced changes in sea level. Prior to the Eocene, transgressions and regressions of the sea were caused by tectonic factors such as changes in the rate of sea-floor spreading, ridge push, thermal subsidence, and sediment/water loading. Since the Eocene, expansion and contraction of continental glaciers have added further complications to sea-level trends. The nature of the basement on the margin also has influenced sea-level trends in the region. On the Atlantic margin, basement consists of a mosaic of allochthonous terranes docked onto the North American craton

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Figure 58. Isopach map, in meters, of postglacial Holocene sediments. Note that the major depocenter of these sediments is in Provincetown Hook. Compiled from O’Hara and Oldale (1987, Fig. 7), Oldale and O’Hara (1980, Fig. 8; 1990, Fig. 9), Oldale et al. (1973, Plate 2B), and seismic reflection profiles recorded by Ballard and Uchupi (1975) and archived at the Woods Hole Oceanographic Institution.

during the closing of the Paleozoic Atlantic (e.g., Uchupi and Aubrey, 1988). Because these terranes varied in their tectonic fabric, they have reacted differently to the above changes so that sea level on the margin has varied not only with time but also along its length. The possibility that the continental plateaus are reacting independently from oceanic basins to

internal forces further complicates the history of sea-level changes. Thus, to ascribe the sea-level trend documented by tide gauges in New England solely to glacial unloading may not be realistic (see Emery and Aubrey, 1991, p. 23–52, and references therein for a more detailed discussion on the causes of relative sea-level changes).

Late Quaternary construction of Cape Cod, Massachusetts ACKNOWLEDGMENTS This synthesis on the geology of Cape Cod and vicinity would not have been possible without numerous discussions we had over the years with R. N. Oldale, who has devoted several decades of his research activities to the geology of the Cape and adjacent offshore area. Funds provided via the J. Seward Johnson Chair in Oceanography awarded to Uchupi by the Woods Hole Oceanographic Institution (WHOI) made possible his contribution to this study. WHOI also provided some funding to complete the graphics. Participation of Giese and Aubrey was made possible by funds provided by the Minerals Management Service of the U.S. Department of the Interior through a cooperative agreement with the Massachusetts Office of Environmental Affairs. The Texas Bureau of Economic Geology and the Massachusetts Office of Coastal Zone Management, through the efforts of J. O’Connell, facilitated the project administration. This work is partly a result of research sponsored by National Oceanographic and Atmospheric Administration (NOAA) National Sea Grant College Program Office, Department of Commerce, under Grants NA90-AA-D-SG480, WHOI Sea Grant Project R/G-20-PD and NA46RG0470, and WHOI Project R/G-22-PD. Comments by R. N. Oldale, M. Allison, and R. S. Williams, Jr., were helpful in the completion of this work. Woods Hole Oceanographic Institution Contribution No. 9041. REFERENCES CITED Aubrey, D. G., and Gaines, A., 1982, Rapid formation and degradation of barrier spits in areas with low rates of littoral drift. Marine Geology, v. 49, p. 257–278. Aubrey, D. G., and Giese, G. S., eds., 1993, Formation and Evolution of Multiple Tidal Inlets, Coastal and Estuarine Studies: American Geophysical Union Monograph Series 44, 235 p. Aubrey, D. G., and Speer, P. E., 1984, Updrift migration of tidal inlets: Journal of Geology, v. 92, p. 631–545. Aubrey, D. G., Twichell, D. C., and Pfirman, S. L., 1982, Holocene sedimentation in the shallow nearshore zone off Nauset Inlet, Cape Cod, Massachusetts: Marine Geology, v. 47, p. 243–250. Ballard, R. D., and Uchupi, E., 1975, Triassic rift structures in Gulf of Maine: American Association of Petroleum Geologists Bulletin, v. 59, p. 1041– 1072. Bard, E., Hamelin, B., Fairbanks, R. G., and Zindler, A., 1990, Calibration of the 14C timescale over the past 30,000 years using the mass spectrometric U-Th ages from Barbados corals: Nature, v. 345, p. 405–410. Belknap, D. F., and Shipp, R. C., 1988, Deglaciation of the Gulf of Maine: seismic stratigraphic evidence for varying ice margins on the Maine inner shelf: Geological Society of America, Abstracts with Programs, v. 20, p. A134. Belknap, D. F., Shipp, R. C., Kelley, J. T., and Schnitker, D., 1989, Depositional sequence modeling of late Quaternary history, west central Maine coast, in Tucker, R., and Marvinney, R. G., eds., Studies in Maine Geology: Quaternary Geology, v. 5, p. 29–46. Belknap, D. F., Schnitker, D., Bacchus, T. S., Batchelor, B. A., and Fader, G. B. J., 1991, New seismic and core data from the Gulf of Maine: glaciomarine and grounding line deposits: Geological Society of America, Northeastern and Southeastern Sections, Abstracts with Programs, v. 23, p. A7.

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